'Tillite', Boston Basin

Home , Conception Group, Gaskiers Formation, Gaskiers glaciation, Squantum

SEDIMENTOLOGY OF THE SQUANTUM ‘TILLITE’, BOSTON BASIN, USA: MODERN ANALOGUES AND IMPLICATIONS FOR THE PALEOCLIMATE DURING THE GASKIERS GLACIATION (c. 580 Ma)

Shannon Leigh Carto

A thesis submitted in conformity with the requirements for the degree of Doctorate of Philosophy Graduate Department of Geology University of Toronto

SEDIMENTOLOGY OF THE SQUANTUM ‘TILLITE’, BOSTON BASIN, USA: MODERN ANALOGUES AND IMPLICATIONS FOR THE PALEOCLIMATE DURING THE GASKIERS GLACIATION (c. 580 Ma)

Shannon Leigh Carto Doctorate of Philosophy, 2011 Graduate Department of Geology University of Toronto

ABSTRACT

The Gaskiers glaciation (c. 580 Ma) has been classically traced along the

Neoproterozoic Avalonian-Cadomian Terranes, which are now found scattered around the

North Atlantic Ocean. Around 625 Ma these terranes were composed of volcanoes and arc- type basins. ‗Till-like‘ diamictite horizons identified within these basins have been used as evidence for a ‗Snowball Earth-type‘ glaciation at 580 Ma. However, others argue that these deposits are non-glacial debris flow deposits. To test the non-glacial interpretation of these deposits, a detailed sedimentological and basin analysis was conducted on the Neoproterozoic

Squantum Member that occurs conformably with the volcanic-sedimentary rocks of the Boston

Bay Group (eastern Massachusetts); this deposit is one of the most referenced ‗tillite‘ deposits for the Gaskiers glaciation. This thesis shows that the ‗tillites‘ of this succession are volcanically-influenced non-glacial debrites. Using the Lesser Antilles Arc and the adjacent

Grenada Basin in the Caribbean Sea as a modern depositional analogue for the Avalonian-

Cadomian Terranes, this study further reveals that debris flow facies types (diamicts) comparable to those of the Avalonian-Cadomian Terranes are produced at this modern arc and are recorded in the fill of the Grenada Basin. A similar study was conducted on the modern diamicts produced at the heavily glaciated Mount Rainier volcano (Washington, USA),

ii revealing that despite the presence of local glaciers, debris flow is the dominant process depositing diamicts due to eruptions and flood events. The major thrust of this thesis is that it highlights the key role of tectonics and volcanism, not glaciation, in producing the diamictites of the Avalonian-Cadomian Terranes, and the importance of examining Neoproterozoic diamictite facies in their wider sedimentary, stratigraphic and tectonic context.

iii

ACKNOWLEDGEMENTS

When I first embarked on this journey almost five years ago the most that I hoped to attain was a Ph.D. Now as my journey comes to an end I realize I am coming away with so much more. In addition to acquiring an invaluable set of skills and a sound knowledge base in paleoclimatology, I have also learned important life lessons, built life-long friendships, and now proudly possess a ‘worn and torn’ passport that bears the memory of the many adventures I had while acquiring my Ph.D.; adventures which took me from one coast of North America to the other and to the tops of fiery volcanoes and icy mountains. Most importantly, I am coming away with the feeling that I have waited patiently for since I sat down with my first chemistry set at age eight that, at last, I am a scientist.

For all these gifts and opportunities that have been bestowed on me, it is a pleasure to thank the many people who made them possible.

Foremost, I would like to express my sincere gratitude to my advisor, Dr. Nick Eyles. Throughout my Ph.D. program he provided me with motivation, immense knowledge, sound advice, good teaching, and lots of great ideas. He challenged me to think harder, to think outside the preverbal box and challenge established dogma in the field of paleoclimatology. I would also like to thank the other members of my Ph.D. committee, Dr. Andrew Miall, Dr. Uli Wortmann, Dr. Mathew Wells and Dr. Emmanuelle Arnaud for their detailed and constructive comments, encouragement and all the support they have provided me with throughout my Ph.D. program.

I would also like to thank my many student colleagues for providing a stimulating and fun environment in which to learn and grow. I am especially grateful to my dear friends, Angela Falcon and Louise Daurio, for taking time out of their lives and busy school schedules to assist me in the field. Come rain, shine, snow or extreme heat they were by my side as we climbed active volcanoes and glaciated mountains, trekked through the notorious woods of ‘Big-foot’, escaped the witches of Salem, and navigated the ‘two-lane’ roads of the Caribbean, all in search of that elusive ‘pay dirt’; they are responsible for the warm feelings and memories that I have from my Ph.D. None of this would have been possible without them and I am immensely proud and grateful to call them my friends. I can only hope that our adventures are not over and that ‘Charlie’s Angels’ will ride again. I would also like to thank Kathy Wallace and Tom Meulendyk. Although, our time together has been short, I am grateful for all the laughs, support and advice that we have been able to share.

I am forever indebted to Mike Doughty and Tom Meulendyk for their assistance with the preparation of the graphics for my thesis. Despite their own rigorous schedules and work commitments they provided me with the highest quality of work and were always willing to work within the constraints of my thesis schedule.

It is difficult to find the words that adequately convey the depth of my gratitude to my friends, who despite the distance of oceans, continents and city blocks, have always provided me with endless support, motivation, enthusiasm and have brought so much joy to my life. I am forever grateful to you all for the shoulders you provided me with to lean on, for listening to my thesis writing ‘woes’ and for all the laughs, girls’ night outs, weekend getaways, and the many pep-talks which helped me make the bad times good and good times better.

Most importantly, I wish to thank my parents, Annette and Colyn Carto and my sister Candice for their endless encouragement and support. I have been blessed to have a sister who is not only a great friend but my biggest fan. I truly believe that without her support and role modeling I would not be where I am today. She is the most disciplined, fiercy intelligent and hard-working woman I know. I am especially grateful to my parents who have always supported me in the choices I have made (whether they liked them or not) and who always encouraged me to go after my dreams, no matter how big, no matter how impossible. I am especially thankful to my mother who spent many hours proof- reading the chapters of my thesis and who, along with my father, was always ready to give me kind words of support and encouragement during difficult times.

Last, but definitely not least, I would be remised to not thank by little buddy Pete who spent many long days and long nights by my side as I typed, read and typed some more. I don’t how I would have gotten through it all without you. Having you in my life is a blessing and a constant joy.

~There is a single light of science, and to brighten it anywhere is to brighten it everywhere~

Isaac Asimov

TABLE OF CONTENTS

ABSTRACT……………………………………………………………………..…...... ii ACKNOWLEDGEMENTS ………………………………………………………….………iv TABLE OF CONTENTS………………………………………………………………...…...vi LIST OF FIGURES ……………………………………………………………………...... ix LIST OF TABLES…………………………………………………………………..….…....xvi CHAPTER 1: INTRODUCTION……………………………………………….……..…1 1.1 STUDY RATIONALE…………………………………………………...... 1 1.2 OBJECTIVES AND STRUCTURE OF THESIS.………………………..……...…...5 REFERENCES………………………………………………………….…………….…...... 10

CHAPTER 2: NON-GLACIAL, DEEP WATER MASS FLOW ORIGIN FOR THE NEOPROTEROZOIC SQUANTUM ‘TILLITE’ (BOSTON BASIN, USA): NO EVIDENCE OF SNOWBALL EARTH DURING THE GASKIERS GLACIATION AT c. 580 Ma...……………….……...... 15

ABSTRACT……………………….………………….………………………………...... 15 2.1 INTRODUCTION AND PURPOSE OF STUDY.….……....……………...... 17 2.2 PHYSICAL SETTING AND STRATIGRAPHY OF THE BOSTON BASIN…………………………..………………….…...... 18 2.3 PREVIOUS INVESTIGATIONS AND AGE OF THE SQUANTUM MEMBER……………………….………………………...... 21 2.4 SCOPE OF STUDY AND METHODS ……………….…………….………...... 23 2.5 DESCRIPTION AND INTERPRETATION OF LITHOFACIES..……...…...... 24 2.5.1 Conglomerate facies………………………………………………..……...... 25 Interpretation………………………..………………..………...... 28 2.5.2 Sandstone facies…………………………………………………..………...... 31 Interpretation………………………………………………...... 32 2.5.3 Diamictite facies…………………………………………………..…...…...... 33 Interpretation…………………………………………………...... 36 2.5.4 Argillite facies……………………………………………………...…...... …..38 Interpretation…………………………………………...... …...... 40 2.5.5 Volcanic-sedimentary facies…..…………………………….…...... 41 2.6 DEPOSITIONAL AND TECTONIC SETTING……..…………………....…...... 43 2.7 DISCUSSION………………………………….…………………………...... 47 2.8 CONCLUSIONS……………...…………………………………………...…...... 54 REFERENCES……………………………………………………………………...... 79

CHAPTER 3: REINTERPRETATION OF ‘ICE-RAFTED’ PEBBLY ARGILLITES OF THE NEOPROTEROZOIC SQUANTUM MEMBER, BOSTON BASIN, USA AS NON-GLACIAL DEBRITE-TURBIDITE COUPLETS ………………………………………………………….…….……….…...... ….97 ABSTRACT………………………………………………….….…………………..….…...... 97 3.1 INTRODUCTION AND PURPOSE OF STUDY ……………………..…..…..….….99 3.2 GEOLOGICAL SETTING, STRATIGRAPHY

AND AGE OF THE BOSTON BASIN.……………………………………...……....101 3.3 FIELD AND LABORATORY METHODS…..…………………….……….…...... 103 3.4 DESCRIPTION OF LAMINATED ARGILLITE FACIES AT THE SQUANTUM HEAD SECTION...... 104 3.5 INTERPRETATION OF FACIES AND RECONSTRUCTION OF DEPOSITIONAL SETTING….....………………...... …107 3.6 DISCUSSION….....……………………………………...……………….….…...…..111 3.7 CONCLUSIONS…………………………………………...……………..…….…....115 REFERENCES ……………………………………………………....….….………...... 141

CHAPTER 4: MASS FLOW REWORKING OF GLACIAL SEDIMENT ON MOUNT RAINIER: IMPLICATIONS FOR UNDERSTANDING THE FATE OF TERRESTRIAL GLACIAL SEDIMENT ON NEOPROTEROZOIC ISLAND ARCS……………………...……..….157

ABSTRACT……………………………………………………………………….…...…...... 157 4.1 INTRODUCTION AND PURPOSE OF STUDY ………………...……………...... 159 4.2 PHYSICAL SETTING AND SEDIMENTARY DEPOSITS OF MOUNT RAINIER ………….…………………………………………...... 161 4.3 STUDY AREA, METHODS AND TERMINOLOGY ……………………....……...164 4.4 STUDY RESULTS……………………..…………………………………...... 165 4.4.1 Glacial till/drift ……………………………………………………….…...... 165 4.4.2 Osceola Mudflow……...……………………………………………...... ….168 4.4.3 Outburst flood deposits…………………………………………....……...…..171 4.5 INTERPRETATION..…....…………………………………….…………..…..….....172 4.6 DISCUSSION……………………………………………………………...... …...175 4.7 CONCLUSIONS.………………..………….…………………..…...….....……..…..175 REFERENCES………………………………………………………………...... …204

CHAPTER 5: VOLCANIC-SEDIMENTARY MASS FLOW FACIES OF THE LESSER ANTILLES ARC AND GRENADA BASIN: MODERN ANALOGUES FOR NEOPROTEROZOIC ‘TILLITES’ OF THE AVALONIAN-CADOMIAN TERRANES………………………………………...……....212

ABSTRACT………………………………………………….………………………..……...212 5.1 INTRODUCTION AND PURPOSE OF STUDY……………………..……....…...... 214 5.2 GEOLOGICAL SETTING AND TECTONIC HISTORY OF THE LESSER ANTILLES ARC……………………………….….....…………..216 5.3 STUDY AREAS AND METHODS……………………………..…………..…...... 218 5.4 DESCRIPTION OF SUBAERIAL MASS FLOW FACIES...... ……...…..221 5.4.1 Debris avalanche deposits………….………………………………...... 221 5.4.2 Lahar deposits……………………….………………….……………...... 223 5.5 OFFSHORE MASS FLOW FACIES IN THE GRENADA BASIN….……………..226 5.6 IMPLICATIONS FOR UNDERSTANDING NEOPROTEROZOIC PALEOCLIMATE DURING THE GASKIERS GLACIATION

vii

AT C. 580 Ma..….…………………………………………………....………...... 231 5.7 CONCLUSIONS……………………………………………………………...…..….235 5.8 REFERENCES…………………….………………………………...……...... …...... 245

CHAPTER 6: SUMMARY ……………………………………….……...….…..….259 6.1 SIGNIFICANCE AND CONCLUSIONS …………………………...….……..…….259 6.2 OPPORTUNITES FOR FURTHER RESEARCH……………………...………...... 262 REFERENCES……………………………………………………………...…………...... 263

APPENDIX 1: Glossary………………………………………………...……………..……265

viii

LIST OF FIGURES

Figure 2.1: Reconstruction of Gondwana at c. 570 Ma showing paleogeographic position of the Avalonian-Cadomian Belt (Neoproterozoic island arc peri-Gondwanan terranes)……………...………...…...... 55

Figure 2.2: (A) Location of the Boston Basin and type locality of the Squantum Member at Squantum Head…………………………………………….…...56

Figure 2.2: (B) Simplified geology of the Boston Basin (outlined in dashed line) showing location of study sites………………..…….………………………………….…...... 57

Figure 2.3: (A) Clast-supported pebble-to-boulder conglomerate. (B) Cobble-to-granule conglomerate partially mixed with. laminated argillite facies showing soft-sediment deformation………………….…………58

Figure 2.3: (C) ‗Close-up‘ of partially mixed cobble-to-granule conglomerate and laminated argillite facies shown in Figure 2.3B. (D) Matrix- supported pebble-cobble conglomerate supported by coarse sand matrix…………………………………………………………………………………...….59

Figure 2.3: (E) Matrix-supported pebble-cobble conglomerate with massive coarse-grained sandstone lenses. (F) Crudely stratified succession of pebble conglomerate with beds of coarse-grained granule-sandstone……………………………….………...... ….....60

Figure 2.4: (A) Massive, fine-grained sandstone with deformed coarse-grained sandstone lenses or ‗wisps‘. (B) Low-angle cross-laminated medium- to fine-grained sandstone……………………………………………….……………...…..61

Figure 2.4: (C) Laminated and normally graded sandstone. (D) A succession of laminated and normally graded sandstone deformed by large overturned fold………...….62

Figure 2.4: (E) Deformed laminated and normally graded sandstone showing pillow structures. (F) Medium-to-fine-grained massive sandstone displaying ‗dish- and-pillar‘ water-escape features…………….…………………………....……….…...….63

Figure 2.5: (A) Matrix-supported, homogeneous diamictite. (B) Clast-supported diamictite showing crude grading…………………………………………………..…..….64

Figure 2.5: (C) Heterogeneous diamictite showing chaotically mixed clast-supported conglomerate with muddy-siltstone. (D) Heterogeneous diamictite showing incomplete mixing of matrix-rich diamictite with clast-rich diamictite…………….……..65

Figure 2.5: (E) Rounded, massive mudstone slab within matrix-supported diamictite. (F) Laminated and normally graded argillite………………….……………….….….…....66

Figure 2.6: (A) Diamictite at type section interbedded with a thin succession (52 cm in total) of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies. The locations of stratigraphic logs shown in Figures 2.8A-C are labeled (logs A-C). (B) Close-up of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies at Squantum Head...... 67

Figure 2.6: (C) Cut-slab of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies exposed at type section………...……...…..68

Figure 2.7: (A) Beds of reworked lapilli tuff within a succession of laminated, normally graded argillite exposed along the southeastern coast of Squantum Head. (B) Massive bed of reworked lapilli tuff (refer to Figure 2.7 A)...... …..69

Figure 2.7: (C) Photomicrograph of lapilli tuff showing subangular to subrounded clasts of quartz, iron oxides, alkali feldspar, and fresh sodic plagioclase, abundant fine felsic (rhyolite) tuffs, and rounded volcanic clasts of andesite and dacite. (D) Photomicrograph of laminated, normally graded argillite consisting of subangular-subrounded silt-sized quartz, plagioclase, feldspar, melaphyre and dark grains of volcanic glass……………………………….…………..……………..……...... 70

Figure 2.8: Stratigraphic logs of key outcrops examined in study area. Logs A-C are from the type site (study site 1; Fig. 2.2B)…………...…………………..………….....71

Figure 2.8: Stratigraphic logs of key outcrops examined in study area. D: site 1, E: site 10 and F: site 10 (Fig. 2.2B). …..………..…………………..…….....…..72

Figure 2.8: Stratigraphic logs of key outcrops examined in study area. G: site 12, H: site 1, and I: site 1 (Fig. 2.2B).……...……………………...……...... 73

Figure 2.9: A 2.5 m-thick clast-supported conglomerate bed resting on (at base of hammer) a pyroclastic flow deposit (Brighton Volcanics)…………...……...... 74

Figure 2.10 Conceptual model of subaqueous debris flow origin for Squantum diamictites and associated turbidite facies involving the episodic downslope slumping and mixing of fluvial conglomerate/gravel detritus with basin slope sand and mud…………………………………….……………...75

Figure 2.11 (A) A thick succession (16.5 m thick) of massive, clast-supported conglomerate conformably overlying heterogeneous diamictites, which in turn conformably overlies heavily-deformed massive argillite (site 1; Fig. 2.2B). (B) Corresponding stratigraphic log of outcrop shown in Figure 2.11A, highlighting the lithologic gradations that exist between conglomerate, diamictites and argillite facies at this section. (C) Heterogeneous diamictites underlying conglomerate units. (D) Contact between heterogeneous diamictites and lower massive and deformed argillite. (E) Lower-most unit of massive and deformed argillite.…..…….....…76

Figure 2.12: Schematic of inferred paleoenvironment of the Boston Basin between c. 595-570 Ma with cross-section through the basin fill showing inferred stratigraphic relationships and depositional mechanisms of the lithofacies of the Boston Bay Group.…………………………………………..…..…..77

Figure 2.13: Map of eastern U.S.A. showing location of the ‗Avalonian Olistostromes‘. (A) Westboro Formation located north and west of the Boston Basin, (B) Blackstone Group in northern Rhode Island, and (C) Newport Neck Formation in southern Rhode Island………..…………….……....78

Figure 3.1: Location of Boston Basin in eastern Massachusetts, USA, showing location of study site of Squantum Member at Squantum Head as well as locations of lithofaices shown in Figure 3.2A-L...... ……..….....117

Figure 3.2: (A) Massive, clast-supported cobble-to-boulder conglomerate. (B) Massive, matrix-supported pebble-cobble conglomerate supported by coarse sand matrix…………………………………………………………………….118

Figure 3.2: (C) Stratified pebble-cobble conglomerate with coarse- grained sandstone. (D) Succession of laminated and normally graded sandstone…………………….……………………………….….……..119 . Figure 3.2: (E) A succession of normally graded sandstone-siltstone laminae deformed by large fold. (F) Massive fine-grained sandstone underlying heavily weathered diamictite bed…………………………………………………....…...120

Figure 3.2: (G) Medium-to fine-grained sandstone showing low-angle cross-lamination. (H) Matrix-supported, homogeneous diamictite…..…….….….…...... 121

Figure 3.2: (I) Clast-supported diamictite showing crude normal grading. (J) Heterogeneous, matrix-supported diamictite…………………………………...... ….122

Figure 3.2: (K) Laminated and normally graded argillite. (L) Bed of reworked lapilli tuff within a succession of laminated and normally graded argillite…….…....…...123

Figure 3.3: Satellite image of Squantum Head showing location of study site.………………....…....124 . Figure 3.4: (A) Study site at Squantum Head where a succession of matrix-supported diamictite is interbedded with a thin succession (52 cm in total) of alternating composite units of laminated pebbly argillite and laminated, normally graded argillite. (B) Close-up of alternating composite units of laminated, normally graded argillite and laminated pebbly argillite………………………………………..…..125

Figure 3.4: (C) Close-up of study site showing two massive diamictite beds interbedded by a thin succession of alternating units of laminated pebbly argillite and laminated and normally graded argillite (heavily eroded area). (D) Stratigraphic log of study section shown in Figure 3.4C. (E) Stratigraphic log of a single composite unit of laminated and normally graded argillite shown in Figure 3.4D. (F) Stratigraphic log of a single composite unit of laminated pebbly argillite shown in Figure 3.4D. (G) Diagram of pebbly argillite layers (shown in Figure 3.4E) showing each unit composed of a lower micro-laminae of diamictite overlain by a massive, micro-laminae of silty mudstone. …….………………….………………………………….….…..….…126

Figure 3.5: (A) Cut slab of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies at Squantum Head…………...….……..…127 . Figure 3.5: (B) Photomicrograph of a thin section of a normally graded argillite lamina showing contact between successive lamina (dashed line). (C) Photomicrograph of a thin section of a pebbly argillite lamina showing contact with overlying mudstone lamina………………………………………..128

Figure 3.5: (D) Cut-slab of laminated pebbly argillites indicating large igneous clasts that are ‗enveloped‘ by underlying and overlying laminae (e.g., clasts I, II, and III). ………………………………..…..………….…….……….….129

Figure 3.6: (A) Succession of laminated, normally graded and laminated pebbly argillites at study site (Fig. 3.4A). (B) Close-up of ‗outsized clasts‘ within laminated pebbly argillites with their long axes parallel to basal surface of lamina (strike NE/ dip 45oSW)……………………..………………………....130

Figure 3.7: (A) Photomicrograph of an argillaceous clast within lower diamictite lamina of a pebbly argillite layer where no clear disruption of sediment around the clast is present. (B) Photomicrograph of a quartzose clast within lower diamictite lamina of a pebbly argillite layer creating a distinct ‗pinch and swell‘ appearance with overlying lamina. Note: no clast disruption occurs at the base of the clast. (C) Photomicrograph of a lithic clast within lower diamictite lamina of a pebbly argillite layer creating a distinct ‗pinch and swell‘ appearance at the upper boundary of the enclosing laminae.……………..………...... 131

Figure 3.8: Thinly-bedded, subaqueous Devonian laminae of basaltic composition deposited during submarine eruptions (Madfeld, Germany)…………….....132

Figure 3.9: (A) Composite unit of pebbly argillite layers measured for ‗maximum clast size versus bed (lamina) thickness‘ analysis. Boundaries of lamina (black lines) and clasts used for measurement (lettered and lined) are shown. (B) Scatter-plot of ‗maximum clast size and laminae (lamina) thickness‘ data for the composite sequence of laminated pebbly argillite shown in Figure 3.9A.………………………….…..……….……………………….…....133

Figure 3.10 Conceptual model of the evolution of ‗co-genetic debrite-turbidite‘ layers (laminated pebbly argillites) exposed at Squantum Head, as described in the text. ………………...……....……………………………………………..……..…134

xii

Figure 3.11: (A) Paleogeographic map of microcontinents at c. 580 Ma showing the distribution of Gaskier-aged glacigenic formations (stars) identified by Hoffman and Li (2009). (B) Reconstruction of microcontinents ~565 Ma showing locations of Gaskiers glacial deposits reported by Kawai et al. (2008)…...... ….135

Figure 4.1: Schematic cross section through the Juan de Fuca Ridge and the Cascadia subduction zone showing location of Mount Rainier within the Cascade Range. ……………………………………………...……….…….....179

Figure 4.2: General schematic of the distribution of glacial drift and outburst flood deposits and present-day glaciers on Mount Rainier showing location of ‗Glacial drift/till‘ and ‗Outburst flood‘ study sites……….………………………………………...…...…...180 . Figure 4.3: Ice-contact columnar jointed andesite lava flows at Mount Rainier (Burroughs Mountain)………….…….………….……….……..181

Figure 4.4: Extent of major lahars generated at Mount Rainier over the past 6000 years and locations of Osceola Mudflow study sites ………………………..…………...182

Figure 4.5: (A) Evans Creek glacial drift (locality C; Fig. 4.2) overlain by a thin bed of the Paradise Lahar. (B) Evans Creek glacial drift (locality D; Fig. 4.2)………………………………………………………………………183

Figure 4.5: (C) Evans Creek glacial drift (locality E; Fig. 4.2). (D) Evans Creek glacial drift (locality F; Fig. 4.2) overlying striated rhyolite……………………….….…184

Figure 4.5: (E) Close-up of Evans Creek glacial drift shown in Figure 4.5D (locality F; Fig. 4.2) (F) Evans Creek glacial drift (locality G; Fig. 4.2) overlying basalt.……………..….………………………………………...….....185

Figure 4.5: (G) Close-up of Evans Creek glacial drift shown in Figure 4.5F (locality G; Fig. 4.2). (H) Evans Creek glacial drift (locality H; Fig. 4.2).…………..….…………………………………………………………....…...... 186

Figure 4.5: (I) Close-up of Evans Creek glacial drift shown in Figure 4.5H (locality H; Fig. 4.2). (J) Close-up of Evans Creek glacial drift shown in Figure 4.5H (locality H; Fig. 4.2)……………..….……….…...….187

Figure 4.6: Scatter diagram of clast roundness vs. elevation data collected from deposits of Evans Creek glacial drift at locations shown in Figure 4.2, showing increased roundness of clasts down valley. ……………………..…188

Figure 4.7: (A) Hayden Creek glacial drift (locality I; Fig.4.2). (B) Close-up of Hayden Creek glacial drift shown in Figure 4.7A (locality I; Fig. 4.2).……….…...….189

Figure 4.8: (A) Vashon Till deposit in the Ohop Valley in the Puget Sound Lowland (locality A; Fig. 4.2). (B) Close-up of Vashon Till shown in Figure 4.8A……………………………..…………………………...…...... ….190

Figure 4.9: Stratigraphic logs of key outcrops of Evans Creek Drift. (A) locality D (Figs. 4.2, 4.5B); (B) locality F (Figs. 4.2, 4.5 D/E)…………………………...……..….191

xiii

Figure 4.9: Stratigraphic logs of key outcrops of (C) Evans Creek Drift (locality G; Figs. 4.2, 4.5F-G) and (D) Vashon Till (locality A; Figs. 4.2, 4.8A/B)…….…...…..…192

Figure 4.10: (A) Osceola Mudflow (outlined in square) exposed on the floor of Glacier Basin at locality K; Fig. 4.4. (B) Striated boulder within the Osceola Mudflow in Glacier Basin (locality K; Fig. 4.4). (C) Striated clasts found in the Osceola Mudflow (locality K; Fig. 4.4)....…..…….….193

Figure 4.11: (A) Osceola Mudflow overlying Evans Creek drift (locality L; Fig. 4.4). (B) Osceola Mudflow exposed along Sunrise Road (locality M; Fig. 4.4).……………………………… ……….……..………..…..…...... 194

Figure 4.11: (C) Osceola Mudflow at locality N (Fig. 4.4). (D) Close-up of Osceola Mudflow shown in Figure 4.11C (locality N; Fig. 4.4)………………...……..195

Figure 4.11: (E) Osceola Mudflow in the town of Greenwater at locality O (Figs. 4.4). (F) Close-up of Osceola Mudflow shown in Figure 4.11E (locality O; Fig. 4.4) composed of beds of massive diamict and massive and crudely laminated fine-grained silty sandstone………………..……..196

Figure 4.11: (G) Close-up of fine-grained crudely laminated silty-sandstone bed of the Osceola Mudflow shown in Figure 4.11F (locality O; Fig. 4.4). (H) Close-up of massive diamict bed of Osceola Mudflow shown in Figure 4.11F (locality O; Fig. 4.4). …………………………….....197

Figure 4.11: (I) Osceola Mudflow at locality O (Fig. 4.4) showing reverse grading defined by upward transition from pebble/cobbles to boulders. (J) Distal bedded Osceola Mudflow deposit at Mud Mountain Dam (locality P; Fig. 4.4).………………………………………….….…….198

Figure 4.12: (A) Scatter diagram of clast roundness vs. elevation data collected from deposits of Osceola Mudflow in White River valley (Fig. 4.4). (B) Scatter diagram of number of striated clasts vs. elevation collected from deposits of Osceola Mudflow in the White River valley (Fig. 4.4)………....…....199

Figure 4.13: Stratigraphic logs of key outcrops of Osceola Mudflow. (A) locality K (Figs. 4.4, 4.11A), (B) locality N (Figs. 4.4, 4.11C/D), and (C) locality O (Figs. 4.4, 4.11 E-I)……………………………………..……...... …200

Figure 4.14: (A) Outburst flood deposit exposed in Kautz Creek (locality Q; Fig. 4.2). This deposit records the 1947 flood event derived from the Kautz Glacier. (B) Tahoma Creek outburst flood deposit recording flood events in the 1960‘s (locality R; Fig. 4.2)…………………………………..…....201

Figure 4.14: (C) Outburst flood deposit exposed in the Van Trump Creek deposited at circa. 2001(locality S; Fig. 4.2)……… …………………………...……....202

Figure 4.15: Stratigraphic log of Kautz Creek Outburst flood deposit (locality Q; Figs. 4.2, 4.14A)…………………………………………...... ………...... 203

xiv

Figure 5.1: (A) General schematic of main components of the Lesser Antilles island arc system located at the eastern margin of the Caribbean Plate………..………237

Figure 5.1: (B) Bathymetry of the eastern and western sides of the Lesser Antilles Arc showing location of study sites on islands included in study, the extent of different debris avalanche deposits on the sea floor, and main active volcanoes on each island…...………….……..………….…238

Figure 5.2: (A) Block-supported debris avalanche deposit (Wotten Waven Sulphur Springs, Dominica; Fig. 5.1B). (B) Block- supported debris avalanche deposit overlying a matrix-supported debris avalanche deposit (Choisel Beach, St. Lucia; Fig. 5.1B)…………...….…...... …239

Figure 5.3: (A) Poorly sorted massive lahar deposit (Wotten Waven Springs, Dominica; Fig. 5.1B). (B) Poorly sorted massive lahar deposit along (Malendure, Basse Terre of Guadeloupe (Fig. 5.1B). (C) Poorly sorted massive lahar deposit (Grande Anse Beach, Basse Terre of Guadeloupe; Fig. 5.1B). (D) Moderately sorted massive lahar deposit overlain by thin bed of crudely laminated fine-grained clay-silt-sand (Grande Anse Beach, Basse Terre of Guadeloupe; Fig. 5.1B). (E) Lahar deposit showing incomplete mixing of finer grained matrix material within coarse-grained sediment (Rosalie Point, Dominica; Fig. 5.1B). (F) Massive, matrix-supported lahar deposit which transitions upward into clast-supported, reversely graded lahar (Grande Caille Point, St. Lucia Fig. 5.1B)………………………… …....….…..……..240

Figure 5.4: Schematic illustration of the asymmetric distribution of volcanogenic deposits east and west of the Lesser Antilles Arc and the factors chiefly responsible for this dispersal pattern. The figure portrays an eruption plume, which is transported to the east by the prevailing westerlies, resulting in air-ash layers in the Atlantic Ocean east of the arc. Contemporaneous debris flows enter the sea west of the arc where the slopes are steep…..………...…………………………………….….…...241

Figure 5.5: Schematic east-west cross-section of the Lesser Antilles arc system showing the distribution of volcanogenic sediments in and around the active arc …………………………………………………...... …..242

Figure 5.6: Stratigraphic logs of subaerial mass flow facies from the Lesser Antilles Arc. (A) Wotten Waven Springs, Dominica (Figs. 5.1B, 5.2A), (B) Choisel Beach, St. Lucia (Figs. 5.1B, 5.2B), and (C) Grande Anse Beach, Basse Terre, Guadeloupe (Figs. 5.1B, 5.3C/D)…………….……………… …………………………………………………..243

Figure 5.7: Lithology of sediment core retrieved from the deepest part of the Grenada Basin, based on deep-sea core named ―CAR-DOM-1‖ (CARAVAL cruise, 2002) located 120 km west of the island of Martinique. Diagram shows interbedding of volcaniclastic ash layers, hemipelagic sediments and thick diamictite deposits………………………...... 244 xv

LIST OF TABLES

Table 1.1: List of diagnostic features used for the recognition of glacial, glacially-influenced, and non-glacial mass flow diamict(ite)s……..…..………...25

Table 3.1: Descriptions and interpretations of lithofacies of the Boston Bay Group …………………………...……………………….…….…..136

Table 3.2: Maximum clast size and laminae thickness data for laminated pebbly argillite facies……………...…………………………..….....138

Table 3.3: Brief description of ice-rafted facies identified within formations assigned to the Gaskiers glaciation (based on those reported by Kawai et al. (2008) and Hoffman and Li (2009)……....…139

xvi

CHAPTER 1:

INTRODUCTION

1.1 STUDY RATIONALE

Paleoclimate models portraying the behavior of the Earth‘s hydrosphere and atmosphere in the remote geologic past are fundamentally reliant on and constrained by data derived from field and laboratory analyses of ancient sedimentary rocks. In regard to

Neoproterozoic climates, the ‗Snowball Earth hypothesis‘ (SEH; Hoffman et al., 1998; Evans,

2000; Hoffman and Schrag, 2000, 2002; Kirschvink, 2002; Schrag et al., 2002; Hoffman, 2005) advocates that there were at least four so-called global-scale ‗hard Snowball‘ glaciations between 770-550 Ma. This idea was evoked to explain the global distribution of ancient poorly sorted sedimentary rocks (diamictites1), at the time interpreted as glacial tillites2. These glaciations are argued to have been extremely cold events unlike any glaciations since, with thick (~400 m) sea ice covering the world‘s oceans leading to the cessation of the world‘s hydrosphere and profoundly affecting the biosphere (Narbonne, 1998; Hoffman and Schrag,

2000, 2002; Narbonne and Gehling, 2003; MacGabhan, 2005; Stern et al., 2006; MacDonald et al., 2010). However, many facets of the SEH are inconsistent with results of sedimentological studies of the Neoproterozoic glacial record which is represented by glacially influenced marine strata (glacigenic diamictites within deep marine turbidites) recording the supply of abundant meltwater and sediment to open water environments akin to well-known glaciomarine settings of the Phanerozoic (Young, 2002; Leather et al., 2002; Eyles and

Januszczak, 2004; Eyles, 2008; Etienne, 2007; Allen and Etienne, 2008; Zhao et al., 2009;

1 Non-genetic term for lithified sedimentary rocks composed of a poorly sorted mixture of gravel- to-boulder-sized clasts in a fine-grained matrix. 2 Lithified diamictites deposited directly under ice-sheets due to the erosion of underlying bedrock. 1

Zhao and Zheng, 2010). Many other supposed ‗tillites‘ classically held to record widespread

‗global glaciation‘ (see Eyles, 2004) are argued instead to be the product of mass flow (debrite3) in tectonically-active deep marine basins exhibiting no definitive evidence of ice (see Table 1.1 for a list of diagnostic features used for the recognition of glacial, glacially influenced and non- glacial debris flow diamictites)(Schermerhorn, 1974, 1983; Eyles, 1993; Crowell, 1999; Eyles and Januszczak, 2004, 2007; Eyles, 2008). The terrestrial record of glaciation is very restricted by comparison, with notable exceptions such as across the North African craton (see Eyles and

Januszczak, 2004 for review). Significantly, Neoproterozoic glaciations occurred in the broader tectonic context of the breakup of Rodinia and the formation of rift basins. Accordingly, it is widely argued that the rock record from the Neoproterozoic points to less severe glacial episodes (Leather et al., 2002; Condon et al., 2002; Olcott, 2005; Corsetti et al., 2006; Eyles et al., 2007; Zhao et al., 2009; Zhao and Zheng, 2010) wherein glaciers were tectonically- controlled and regionally-restricted, nucleating on tectonically-generated high topography (e.g., volcanic arcs, uplifted rift margins and orogenic belts; Young, 1995; Eyles and Januszczak,

2004, Eyles, 2004, 2008).

The Gaskiers glaciation at c. 580 Ma (Bowring et al., 2003) is the youngest and shortest of four putative globally cold Snowball Earth events (Hoffman et al., 1998; Hoffman and

Schrag, 2002; Halverson et al., 2005; Hoffman, 2005). The global type deposit (the Gaskiers

Formation of the Avalon Peninsula in Newfoundland) consists of a thick fill of turbidites, volcanic ash and flows, and debrites. The Gaskiers Formation appears to have accumulated in a rift basin within the Avalon Terrane. A glacial influence on sedimentation is present in the form of striated and glacially shaped (faceted and/or bullet-shaped) dropstones within the

3 Lithified diamictites deposited by debris flows, unrelated to climate. 2 succession (Bruckner and Anderson, 1971; Williams and King, 1975; Eyles and Eyles, 1989) recording floating ice released from glaciers growing along the terrane fringed by deep water

(Eyles and Eyles, 1989; Eyles, 1990). The broader setting of the Avalon Terrane is that of interlinked island-arc terranes and marginal arc basins that formed peripheral to the active northern margin of Gondwana in the Late Proterozoic (c. 620-570 Ma) (known collectively as the Avalonian-Cadomian Terranes or peri-Gondwanan terranes), now found scattered around the margins of the North Atlantic Ocean as a consequence of the breakup of Pangea (including

New England, Atlantic Canada, southern Britain, northwestern France and Bohemia; Nance,

1990; Murphy et al., 1999; Murphy et al., 2004). Correspondingly, other broadly correlative arc-related deep marine diamictite-bearing successions occur within the Avalonian-Cadomian

Terranes, where classic interpretations of a glacial influence on diamictite sedimentation held since 1914 are reliant on the simple ‗tillite-like‘ appearance of the diamictites (see Hambrey and Harland, 1981 for review). Despite many detailed sedimentological studies that indicate an entirely non-glacial mass flow origin for these deposits (Crowell, 1957; Schermerhorn, 1966,

1974; Winterer, 1963 Eyles and Eyles, 1989; Eyles, 1990) they continue to be linked uncritically to a notional global or near-global glaciation event at c. 580 Ma (e.g., Kawai et al.,

2008; Hoffman and Li, 2009).

The oldest recognized, perhaps most well known (and controversial) Gaskiers-age

‗glacial‘ deposit is the Squantum Member (Billings, 1976) of the Boston Basin in eastern

Massachusetts, USA (c. 597-570 Ma; Thompson and Bowring, 2000; Ault, 2004). This diamictite deposit is inferred to be a glacial tillite, once again on the basis of the similarity in appearance of the diamictites and laminated argillites to modern day tills and glaciolacustrine varves, respectively (Sayles and LaForge, 1910; Sayles, 1914, 1919; Emerson, 1917; Cameron

3 and Jeanne, 1976; Rehmer and Roy, 1976). Crowell (1957) and Dott (1961) argued for a non- glacial origin; but an overall glaciomarine setting is still evoked on the basis of the tillite-like appearance of the diamictites and a single rare horizon of laminated pebbly argillite long interpreted as ‗ice-rafted‘ debris (Socci and Smith, 1987) (see Carto and Eyles, 2011a for historical review of the Squantum Member).

The long-standing disagreement over the true origin of the diamictites of the peri-

Gondwanan terranes clearly demonstrates the need for a more precise understanding of the depositional environments and processes that led to their evolution and deposition. Despite the fact that Neoproterozoic diamictites are at the center of paleoenvironmental interpretations of the Neoproterozoic, to date, detailed sedimentological studies of these deposits are limited.

Instead, paleoclimatic reconstructions are based on simple stratigraphic descriptions and an over-reliance on the marine isotope record, such as strong shifts in 13C/12C, 34S/32S and

87Sr/86Sr isotopes thought to record drastic changes in ocean circulation caused by glacial cover

(i.e., Hoffman and Schrag, 2002). A better understanding of the paleoclimatic significance of these rocks is not only crucial to our ability to accurately interpret the nature and extent of glaciation and paleoenvironmental change during the Gaskiers glaciation, but also the assessment of the precise nature of climatic influences on the evolution of the biosphere given that the Gaskiers glaciation has been linked with the evolution of ‗complex‘ multi-cellular organisms after 545 Ma (‗Cambrian explosion of multi-cellular life‘; see review in Narbonne,

1998; Macdonald et al., 2010).

1.2 OBJECTIVES AND STRUCTURE OF THESIS

This thesis consists of four principal chapters (2, 3, 4, 5) reflecting the sub-objectives described below. Each chapter is written in the format of a self-contained journal paper and it is intended that each be published separately. This approach necessarily creates some overlap and duplication in introductory material.

The overarching objective of this thesis is to conduct a detailed sedimentological and basin analysis of the Squantum Member and associated strata in order to resolve the nature of its depositional environment and thus, the paleoclimatic significance of this diamictite deposit.

Chapter 2 of this thesis presents newly collected sedimentological and petrographic data for the Squantum Member and associated lithofacies, and uses this data to establish a depositional model for the Boston Basin. Recent work on the structure and stratigraphy of the basin and the broader plate tectonic setting is employed to present a coherent overview of sedimentation in a deep volcanically influenced and tectonically active arc-type basin. The data presented here reaffirms the importance of mass flow (not glacial) activity and the reworking of volcanic debris in creating thick debrites within deep marine turbidites.

A second objective is to conduct a detailed sedimentary analysis of previously identified ‗ice-rafted facies‘ (laminated pebbly argillites) found within the Squantum succession. These results are presented in Chapter 3 and confirm a non-glacial origin for the rocks by presenting evidence that they represent ‗co-genetic debrite-turbidite couplets.‘

The results of field investigations presented in the chapters outlined above provide no support for a glacial influence on sedimentation in the Boston Basin. The dominant influence

5 was the availability of large volumes of volcaniclastic and other debris to a rapidly subsiding fault-bounded and tectonically active arc basin.

During the Gaskiers glaciation, glaciers are envisaged as having been local and restricted to the summit of the highest volcanic cones along the arc. Striated and glacially- faceted dropstones within the Gaskiers Formation record the presence of ice as floating ice tongues over the basin; a terrestrial glacial record has not been found anywhere along the

Avalonian-Cadomian Terranes. To better understand the lack of preservation of a terrestrial record of glaciation along the Avalonian-Cadomian Terranes, in regard to the Gaskiers

Formation, a facies investigation was also conducted at the modern-day heavily-glaciated

Mount Rainier volcano located in Washington State, USA. Data from Mount Rainier show that primary glacial sediments and landforms have a very low preservation potential as a result of steep slopes, frequent earthquakes, high rain-fall, rapid spring snowmelt and above all melting of snow and ice during eruptions that create violent outburst floods, which continuously rework and mix the glacial sediment with volcanic and other terrigeneous debris down the volcano‘s valleys. As discussed in Chapter 4, the surficial geology of Mount Rainier is dominated by debrites (lahars); most significantly, deposits formerly mapped as till sheets have been recently reinterpreted as lahars (e.g., Osceola Mudflow). Mount Rainier provides an excellent modern depositional analogue for Neoproterozoic arc-related glaciations by demonstrating how glacial sediment is destroyed in these volcanic environments due to the mass flow processes that dominate in such settings, leaving no terrestrial record of glaciation.

An excellent depositional analogue for the Boston Basin, as well as the other arc- related diamictite-bearing successions of the Avalonian-Cadomian Terranes, can be found today at the modern island-arc system of Lesser Antilles Arc and the Grenada Basin of the

Caribbean Sea. As such, field data was collected from outcrops of volcaniclastic mass flow facies (debris avalanche and lahar deposits) outcropping on four islands of the Lesser Antilles

Arc (St. Lucia, Dominica, Martinique, and Guadeloupe). This information is presented in

Chapter 5 and complemented by published data of deep sea cores and marine geophysical surveys of the submarine equivalents of on-land mass flow facies (volcaniclastic debrites and turbidites) present in the thick (up to 7 km) deep marine fill of the adjacent Grenada Basin

(Sigurdsson et al., 1980; Carey and Sigurdsson, 1984; Deplus et al., 2001). Data presented in this thesis reveals that these facies types are directly analogous to those identified in the Boston

Basin and the facies types described from other related Neoproterozoic arc basins of the

Avalonian-Cadomian Terranes. The data from the Caribbean presents a new and alternative perspective on the depositional origin of the diamictites of the Avalonian-Cadomian Terranes that is unrelated to climate and underscores the absence of any widespread glacial influence on sedimentation other than in the Gaskiers Formation of Newfoundland.

Finally, Chapter 6 briefly summarizes the results of the individual chapters and integrates the conclusions from each. It identifies the broader implications of the results for understanding the nature of glaciation and environmental change during the Gaskiers glaciation.

The primary conclusion is that the SEH does not provide a framework for understanding deposits of the Avalonian-Cadomian Terranes; the model of ‗catastrophic‘ global cooling is rejected in favor of an uniformitarian approach using modern analogues from the Caribbean and Mount Rainier. With the exception of the Gaskiers Formation, the diamictites of the

Avalonian-Cadomian Terranes bear no evidence of having a glacial source; instead they closely resemble classic volcaniclastic mass flow facies produced in modern arc-related basins.

Possibly ice covers were present in other basins of the Avalonian-Cadomian Terranes but their

7 presence is unrecorded indicating only limited extent of glaciation during Gaskiers time (c.580

Ma).

Table 1.1: List of diagnostic features used for the recognition of glacial, glacially influenced and non-glacial debris flow diamictites.

SUBGLACIAL DEBRITE DIAMICTITE LAHARIC DIAMICTITE RAIN-OUT DIAMICTITE DIAMICTITE ORIGIN Diamict(ites) laid down directly by glaciers Diamict(ite)s deposited by debris flows Landslides or debris flows that are Diamict(ite)s formed as icebergs melt in (primary till deposits). (sediment-gravity flows of large clasts associated with volcanoes. subaqueous environments (lakes and oceans) and supported and carried by a mud-water drop their sediment load. mixture). COMPOSITION Gravel- to boulder-sized clasts set in a Gravel- to boulder-sized clasts set in a Gravel- to boulder-sized clasts set in a Gravel- to boulder-sized clasts set in a matrix of matrix of fine-grained material. Admixture matrix of fine-grained material. matrix of fine-grained material. fine-grained material. Mud with clustered or isolated of local and exotic, heterolithic clasts. Locally-derived heterolithic clasts. Commonly composed of 100 % locally ice-rafted debris (boulder- to gravel-sized stones). derived volcanic debris. May contain petrified wood and other terrigeneous material. SORTING Poorly sorted Poorly sorted. Poorly sorted. Poor to moderate sorting GRADING Absent Commonly shows normal or ‗coarse- Reverse or normal grading is common; Absent or occasionally normally graded. tail‘ grading but can be absent. may be ungraded. BED THICKNESS Variable thickness (up to tens of metres). Commonly very thick (up to several Commonly thick (up to hundreds of Commonly thick, ranging from tens to hundreds of kilometers). metres). metres. BEDDING Bedding is poor or absent with little to no Well-defined bedding. Crude internal bedding. Commonly Laterally continuous sheet-like or channelized bed stratification; slumped, loaded, sheared and May show crude stratification and unstratified. geometries that can show loading, folding, iceberg folded units are common. channeled bedding. Deformed and dump and scour features (‗ice keel turbate features). folded bedding is common. Irregular bed geometries are also common. Typically unstratified. CLAST SHAPE Angular to subrounded. Faceted, striated, Subangular to subrounded. Rounded Commonly subangular to subrounded. Angular to subrounded. Faceted, striated, bullet- bullet-shaped and chatter-marked clasts are clasts are common. Faceted and striated Faceted and striated clasts may be shaped and chatter-marked clasts are common. common. clasts may be present. Striations will be present. multi-directional. CLAST FABRIC Strong fabric in which long axis of clasts Commonly displays a random clast Random clast fabric orientation; can Random clast fabric are aligned parallel to flow with secondary fabric but clasts can show a weak fabric show a weak fabric in which long axis transverse to flow fabric. in which long axis of clasts are parallel of clasts are parallel to flow/slope. to flow/slope. LOWER BED Erosional, deformed or sharp planar Gradational or sharp upper/lower Gradational or sharp, non-erosive. Gradational or sharp erosional unconformities. BOUNDARIES contacts; characterized by disconformities contacts are common. Can be erosive or or unconformities. Commonly lies on deformed and loaded. striated bedrock (two or more striation directions). DISTRIBUTION Cover plains and fill valleys; many mantle Cover plains; mantle all surfaces. Fill Spread onto flat piedmont surfaces and Fill basins forming fan-apron lobes. all surfaces. valleys and basins forming fan, apron fill valleys and basins forming fan- and delta lobes. apron lobes. Predominantly subglacial till sheets Sediment-gravity diamict(ites), slide Lahars, debris avalanche (‗mega- Rain-out diamict(ite)s intimately associated with FACIES associated with glaciofluvial and slump deposits, mass flow block‘) deposits, pyroclastic flow and boulder outwash (ice-contact) and muddy sand, ASSOCIATIONS conglomerates/ gravels (subangular to conglomerates composed of poor to surge deposits, tephra layers (ash and channelized massive-graded gravels and sands, subrounded grains and clasts that are moderately sorted, well-rounded clasts lapilli tuff layers), and lava flows. In sediment-gravity diamict(ites), and laterally strongly imbricated, striated and faceted). and grains. In subaqueous subaqueous environments, these facies persistent massive-laminated muds with dropstone Superimposed gravel bars, braided fluvial environments, these facies are are commonly associated with turbidites horizons (ice-distal). Slump and fold structures are units, glaciofluvial sandy reworked deposits commonly associated with turbidites (thinly- to thickly-bedded massive, common. Push-morainal bank and glaciotectonized (cross-stratified sands and thinly bedded (thinly- to thickly-bedded massive, stratified and graded marine sediments or bedrock at ice-sheet grounding silty shales), sediment gravity flows of graded and cross-stratified sandstones) coarse/medium/fine sandstone, line. Stratified and laminated diamict(ites) and reworked till, and glaciotectonized and massive and laminated muds with laminated siltstone and claystone), ash- blankets of laminated muds with dropstones bedrock/sediment. or without isolated and/or clustered fall layers. dominate in distal glaciomarine environments. outsized clasts.

REFERENCES

Allen, P. A., & Etienne, J.L. (2008). Sedimentary challenge to Snowball Earth. Nature Geoscience, 1, 817.

Billings, M.P. (1976). Geology of the Boston Basin. In: J.W. Skehan & D.P. Murray (Eds.), Studies in New England Geology. Geological Society of America, Memoir, 46, 5-30.

Bowring, S.A., Myrow, P.M., Landing, E., & Ramezani, J. (2003). Geochronological constraints on terminal Neoproterozoic events and the rise of metazoans. NASA Astrobiology Institute General Meeting Abstracts, 113–114.

Bruckner, W.D. & Anderson, M.M. (1971). Late Precambrian glacial deposits in southeastern Newfoundland-a preliminary note. Geological Association of Canada Proceedings, 24, 95-102.

Cameron, B., & Jeanne, R.A. (1976). New evidence for glaciation during deposition of the Boston Bay Group. In: B. Cameron (Ed.), Geology of southeastern New England. New England Intercollegiate Geological Conference, 68th Annual Meeting, 117-134.

Carey, S., & Sigurdsson, H. (1984). A model of volcanogenic sedimentation in marginal basins. Geological Society of London, Special Publications, 16, 37-58.

Carto, S. L. & Eyles, N. (2011a). The Squantum Member of the Boston Basin, Massachusetts, USA. In: E. Arnaud, G. P. Halverson, & G.A. Shields-Zhou (Eds.), The Geological Record of Neoproterozoic glaciations. Geological Society of London, Memoir, (in press).

Carto, S.L., & Eyles, N. (2011b). The Gaskiers Formation, eastern Newfoundland, Canada. In: E. Arnaud, G. P. Halverson, & G.A. Shields-Zhou (Eds.), The Geological Record of Neoproterozoic glaciations, Geological Society of London, Memoir, (in press).

Condon, D.J.; Prave, A.R., & Benn, D.I. (2002). Neoproterozoic glacial-rainout intervals: Observations and implications. Geology, 30(1), 35–38.

Corsetti, F.A, Olcott, A.N., & Bakermans, C. (2006). The biotic response to Neoproterozoic snowball Earth. Palaeogeography, Palaeoclimatology, Palaeoecology, 232, 114-130.

Crowell, J. C. (1957). Origin of pebbly mudstone. Geological Society of America, Bulletin, 68, 993-1010.

Crowell, J.C. (1999). Pre-Mesozoic ice ages: their bearing on understanding the climate system. Memoir: Geological Society of America, 192, 106 pp. Deplus, C., LeFriant, A., Boudon, G., Komorowski, J.C., Villemant, B., Harfod, C., Segoufin, J. & Cheminee, J.L. (2001). Submarine evidence for large-scale debris avalanches in the Lesser Antilles arc. Earth and Planetary Sciences Letters, 192, 145–157.

Dott, R.H. Jr. (1961). Squantum ―tillite‖, Massachusetts—evidence of glaciation or subaqueous mass movements? Geological Society of America, Bulletin, 72, 1289-1305.

Emerson, B.K. (1917). Geology of Massachusetts and Rhode Island. US Geological Survey, Bulletin, 597, 289 pp.

Etienne, J.L., Allen, P.A., Rieu, R., & Le Guerroué, E. (2007). Neoproterozoic glaciated basins: A critical review of the Snowball Earth hypothesis by comparison with Phanerozoic glaciations. In: M. Hambrey, P. Christoffersen, N. Glasser, & B. Hubbard (Eds.), Glacial Sedimentary Processes and Products., IAS Special Publication, 39, 343–399.

Evans, A.D. (2000). Stratigraphic, geochronological, and palaeomagnetic constraints upon the Neoproterozoic climatic paradox. American Journal of Science, 300(5), 347–433.

Eyles, N., & Eyles, C.H. (1989). Glacially influenced deep-marine sedimentation of the Late Precambrian Gaskiers Formation, Newfoundland, Canada. Sedimentology, 36, 601-620

Eyles, N. (1990). Marine debris flows: Late Precambrian ‗‗tillites‘‘ of the Avalonian– Cadomian orogenic belt. Palaeogeography, Palaeoclimatology, Palaeoecology, 79, 73–98.

Eyles, N. (1993). Earth‘s glacial record and its tectonic setting. Earth-Science Reviews, 35, 1 – 248.

Eyles, N. (2004). Frozen in time: concepts of global glaciation from die Eiszeit (1837) to Snowball Earth (1998). Geoscience Canada, 31, 157-166.

Eyles, N., & Januszczak, N. (2004). ‗Zipper-rift‘: A tectonic model for Neoproterozoic glaciations during the break-up of Rodinia after 750 Ma. Earth Science Reviews, 65, 1-73.

Eyles, C.H., Eyles, N., & Grey, K. (2007) Palaeoclimate implications from deep drilling of Neoproterozoic strata in the Officer Basin and Adelaide Rift Complex of Australia: a marine record of wet-based glaciers. Palaeogeography, Palaeoecology, Palaeoclimatology 248, 291- 312

Eyles, N. & Januszczak, N. (2007). Syntectonic sub-aqueous mass flows of the Neoproterozoic Otavi Group, Namibia: where is the evidence for global glaciation? Basin Research, 19, 179– 198.

Eyles, N. (2008). Glacioepochs and the supercontinent cycle after ~3.0 Ga: tectonic boundary conditions for global cooling. Palaeogeography, Palaeoecology, Palaeoclimatology, 258, 89- 129.

Halverson, G., Hoffman, P.F., Schrag, D.P., Maloof, A.C., & Rice, A.H.N. (2005). Toward a Neoproterozoic composite carbon isotope record. Geological Society of America, Bulletin, 117, 1181–1207.

Hambrey, M.J. & Harland, W.B. (1981). Earth‘s Pre-Pleistocene Glacial Record. Cambridge University Press, London, 1004 p.

Hoffman, P.F., Kaufman, A.J., Halverson, G.P., & Schrag, D.P. (1998). A Neoproterozoic Snowball Earth. Science, 281, 1342–1346.

Hoffman, P.F., & Schrag, D.P. (2000). Snowball Earth. Scientific American, 282, 62–75.

Hoffman, P. F., & Schrag D. P. (2002). The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14(3), 129-155.

Hoffman, P.F. (2005). On Cryogenian (Neoproterozoic) ice-sheet dynamics and the limitations of the glacial sedimentary record. South African Journal of Geology, 108, 557–577.

Hoffman, P.F., & Li, Z.X. (2009). A palaeogeographic context for Neoproterozoic glaciation Palaeogeography, Palaeoclimatology, Palaeoecology, 277, 158–172.

Kawai, T., Windley, B.F., Terabayashi, M., Yamamoto, H., Isozaki, Y., & Maruyama, S. (2008). Neoproterozoic glaciation in the mid-oceanic realm: An example from hemi-pelagic mudstones on Llanddwyn Island, Anglesey, UK. Gondwana Research, 14, 105-114.

Kirschvink, J.L. (2002). When all of the oceans were frozen. Recherche, 355, 26–30.

Leather, J., Allen, P., Brasier, M.D., & Cozzi, A. (2002). Neoproterozoic Snowball Earth under scrutiny: evidence from the Fiq glaciation of Oman. Geology, 30, 891– 894.

Macdonald, F.A. Schmitz, M.D., Crowley, J. L., Roots, C.F., Jones, D.S. Maloof, A.C., Strauss, J.V., Cohen, P. A., Johnston, D.T., & Schrag, D.P. (2010). Calibrating the Cryogenian. Science, 327, 1241-1243.

MacGabhann, B. A. (2005). Snowball Earth, the Dalradian Supergroup, and the Cryogenian- Ediacaran Boundary. 48th Irish Geological Research Meeting Programme and Abstracts 33

Murphy, J.B., van Staal, C.R., & Keppie, J.D. (1999). Middle to Late Paleozoic Acadian Orogeny in the northern Appalachians: A Laramidestyle plume-modified orogeny. Geology, 27, 653-656.

Murphy, J.B., Fernandez-Suarez, J., Keppie, J.D., & Jeffries, T.E. (2004). Contiguous rather than discrete Paleozoic histories for the Avalon and Meguma terranes based on detrital zircon data: Geology, 32, 585-588.

Narbonne, G.M. (1998). The Ediacara biota: A terminal Proterozoic experiment in the evolution of life. GSA Today, 8(2), 1-6.

Narbonne, G.M., & Gehling, J.G. (2003). Life after snowball: the oldest complex Ediacaran fossils. Geology, 31, 27-30.

O‘Brien, S.J., O‘Brien, B.H., Dunning, G.R., & Tucker, R.D. (1996). Late Proterozoic and related peri-Gondwanan rocks of the Newfoundland Appalachians. In: R.D. Nance & R.D. Thompson (Eds.), Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado. Geological Society of America Special Paper, 304, 9-28.

Olcott, A.N., Corsetti, F.A., & Sessions, A.L. (2005). Biomarkers of a Neoproterozoic Glaciation, Earth System Processes II.

Picard, M., Schneider, J-C., & Boudon, G. (2006). Contrasting sedimentary processes along a convergent margin: the Lesser Antilles arc system. Geo-Marine Letters, 26 (6), 397-410.

Rehmer, J.A., & Roy, D.C. (1976). The Boston Bay Group: the boulder bed problem, In: B. Cameron (Eds.), Geology of southeastern New England: Princeton, New Jersey, New England Intercollegiate Geologic Conference, 68th Annual Meeting, Guidebook, 71-91.

Sayles, R.W., & LaForge, L. (1910). The glacial origin of the Roxbury Conglomerate. Science, 32, 723-724.

Sayles, R.W. (1914). The Squantum Tillite. Bulletin Harvard Museum Comparative Zoology, 66, 141-175.

Sayles, R.W. (1919). Seasonal deposition in the aqueo-glacial sediments. Memoirs of Museum Comparative Zoology, Harvard College, 47, 67 pp.

Schermerhorn, L.J.G. (1966). Terminology of mixed coarse-fine sediments. Journal of Sedimentary Petrology, 36, 831– 836.

Schermerhorn, L. J. G. (1974). Late Precambrian mixtites: Glacial and/or non-glacial? American Journal of Science, 274, 673–824.

Schermerhorn, L. J. G. (1983). Proterozoic glaciation in the light of CO2 depletion in the atmosphere. Memoirs: Geological Society of America, 161, 309– 315.

Schrag, D.P., Berner, R.A., Hoffman, P.F., & Halverson, G.P. (2002). On the initiation of a snowball Earth. Geochemical Geophysics Geosystems, 3, 1029-1036.

Sigurdsson, H., Sparks, R.S.J., Carey, S.N., & Huang, T.C. (1980).Volcanogenic sedimentation in the Lesser Antilles. Journal of Geology, 88, 523-540.

Socci, A.D., & Smith, G.W. (1987). Recent sedimentological interpretations in the Avalon terrane of the Boston Basin, Massachusetts: Maritime Sediments and Atlantic. Geology, 23, 13-39.

Stern, R.J., Avigad, D., Miller, N.R., & Beyth, M. (2006). Geological Society of Africa Presidential Review: Evidence for the Snowball Earth Hypothesis in the Arabian-Nubian Shield and the East African Orogen. Journal of African Earth Sciences, 44, 1–20.

Williams, H., & King, A.F. (1975). Southern Avalon, Newfoundland. Trespassey map area (IK). Papers Geological survey of Canada, 75(1), 11-15.

Winterer, E. (1963). Late Precambrian pebbly mudstone in Normandy, France. In: A.E.M. Nairn (Ed.), Problems in Palaeoclimatology, 159-178. Interscience, London

Young, G. M. (1995). Are Neoproterozoic glacial deposits preserved on the margins of Laurentia related to the fragmentation of two supercontinents? Geology, 23, 153–156.

Young, G.M. (2002). Stratigraphic and tectonic settings of Proterozoic glaciogenic rocks and banded iron-formations: relevance to the snowball Earth debate. Journal of African Earth Science, 35, 451–466.

Zhao Y.Y., Zheng Y.F., & Chen F.K. (2009). Trace element and strontium isotope constraints on sedimentary environment of Ediacaran carbonates in southern Anhui, South China. Chemical Geology, 265, 345-362.

Zhao Y. Y., & Zheng Y.F. (2010). Stable isotope evidence for involvement of de-glacial meltwater in Ediacaran carbonates in South China. Chemical Geology, 271, 86-100.

CHAPTER 2:

NON-GLACIAL DEEP WATER MASS FLOW ORIGIN FOR THE NEOPROTEROZOIC SQUANTUM ‘TILLITE’ (BOSTON BASIN, USA): NO EVIDENCE OF SNOWBALL EARTH DURING THE GASKIERS GLACIATION AT C. 580 MA

ABSTRACT

Did climates of the Neoproterozic (c. 800-570 Ma) include repeated oscillations of severe global cold and brutally hot interglacials as proposed by the ‗Snowball Earth‘ hypothesis? The answer is fundamentally reliant on detailed assessment of the paleoenvironmental setting of sedimentary successions containing poorly sorted sedimentary rocks called diamictites long interpreted as glacial tillites due to their outcrop resemblance to modern glacial till deposits. The Squantum ‗Tillite‘ (commonly referred to as the Squantum

Member; c. 593-570 Ma) of the ~7 km thick turbidite-dominated Boston Bay Group in eastern

Massachusetts, USA, has a century-long history of study most of it predating acceptance of plate tectonics and sedimentary facies analysis. It is varyingly interpreted as being glacial in origin (deposited in either ice-contact terrestrial or glaciomarine settings) or the product of non-glacial mass flow in a deep marine basin. It is currently interpreted in some quarters (in the absence of any further detailed facies study) as evidence of a severe Ediacaran-aged ‗Gaskiers glaciation‘ at c. 580 Ma. A detailed facies analysis of the Squantum Member and associated facies presented here, stressing the importance of the stratigraphic context of diamictites and their wider basinal and tectonic setting, identifies a lack of unique indicators of a glacially influenced setting. Instead, this study identifies a significant volcanic influence on sedimentation in the form of hitherto unrecognized beds of reworked volcanic lapilli tuff and

15 thick successions of turbidites consisting of reworked ash. Facies types within the Boston Bay

Group are shown to be of a subaqueous sediment-gravity flow origin consisting of clast- and matrix-supported, massive, deformed and stratified conglomerates, massive, deformed and normally graded sandstones, and laminated argillites. Massive (homogeneous) and crudely mixed (heterogeneous) diamictites are debrites produced by the mixing of gravel and mud during subaqueous slumping. Thin horizons of laminated pebbly argillites containing ‗outsized clasts‘ that occur with laminated and graded argillites, classically interpreted as ‗ice rafted facies‘, are shown to be thin debrites within turbidites respectively. The depositional model presented here emphasizes a deep marine submarine slope/fan setting in a mid-latitude volcanic arc basin receiving volumes of debris from a volcanic hinterland. There is no case to be made for the continued usage of the term ‗Squantum Tillite‘ or reference to this deposit as evidence for a hyper-cold ‗Gaskiers glaciation.‘

2.1 INTRODUCTION AND PURPOSE OF STUDY

The Late Neoproterozoic Era (c. 800–550 Ma) saw major changes in the Earth‘s geography involving the protracted breakup of the supercontinent Rodinia and the assembly of

Gondwana (Meert, 2001; Powell and Pisarevsky, 2002; Torsvik, 2003; Meert, 2003). This was accompanied by episodic glaciation (see Fairchild and Kennedy, 2007 for review) and the first appearance of complex, multi-cellular animals (Narbonne, 1998; Knoll et al., 2006; Butterfield,

2007). A key uncertainty is the nature and variability of Neoproterozoic climates which have been interpreted by some as being characterized by repeated episodes (some suggest four events; Hoffman and Schrag, 2000, 2002; MacGabhan, 2005; Stern et al., 2006; MacDonald et al., 2010) of the most extreme glacial and interglacial conditions and corresponding large magnitude glacioeustatic sea level fluctuations so far identified on Earth (Snowball Earth

Hypothesis; SEH; Hoffman et al., 1998; Evans, 2000; Hoffman & Schrag, 2002; Kirschvink,

2002; Schrag et al., 2002; Rice et al., 2003; Hoffman, 2005). This hypothesis has triggered much debate. Are Neoproterozoic diamictites true glacial tillites and a record of severe global glacial events? Many sedimentological studies past and present have identified the importance of subaqueous debris flow processes in generating diamictites in incipient rift basins (Crowell,

1957; Dott, 1961; Schermerhorn, 1974, 1983; Eyles and Eyles, 1983; Eyles, 1993; Crowell,

1999; Einsele, 2000; Young, 2002; Leather et al., 2002; Condon et al., 2002; Eyles and

Januszczak, 2004; Etienne, 2007; Allen and Etienne, 2008; Eyles, 2008). The successions that do record a glacial setting are marine deposits similar to well-known ‗temperate‘ Phanerozoic glaciations characterized by abundant meltwater and open marine conditions (Eyles, 1993;

Leather et al., 2002; Condon et al., 2002; Olcott et al., 2005; Corsetti et al., 2006; Zhao et al.,

2009; Zhao and Zheng, 2010). Lack of agreement, however, between these opposing views is a major impediment to global climate modelling (Evans et al., 1997; Hyde et al., 2000; Crowley et al., 2001; Donnadieu et al., 2003; Peltier, 2004; Pierrehumbert, 2004) and the assessment of any climatic influence on the evolution of the biosphere (Macdonald et al., 2010).

The debate with regard to Neoproterozoic climates, underscores the need for a detailed investigation of the sedimentology, stratigraphy and basinal setting of Neoproterozoic diamictite successions. This thesis presents such a study from eastern North America where a

Neoproterozoic diamictite deposit, known as the Squantum Member, is preserved within the volcanic-sedimentary rocks of the Boston Bay Group of the Boston Basin located in eastern

Massachusetts, USA (Billings, 1976, 1979). The Squantum Member has long been identified as a key North American Neoproterozoic ‗tillite‘ (the ‗Squantum Tillite‘; Sayles, 1914) and is now included as key evidence of a global or near-global glacial event known as the Gaskiers glaciation at c. 580 Ma (Thompson and Bowring, 2000; Narbonne and Gehling, 2003; Bingen et al. 2005; Halverson et al., 2005; Kawai et al., 2008; Hoffman and Li, 2009).

2.2 PHYSICAL SETTING AND STRATIGRAPHY OF THE BOSTON BASIN

The Boston Bay Group is a ~7 km thick marine volcanic-sedimentary succession preserved in the Boston Basin within the Avalon Terrane of eastern Massachusetts, USA.

During the Late Neoproterozoic (c. 640-570 Ma) the Avalon Terrane was linked with several other terranes, which together formed the Avalonian-Cadomian Orogenic Belt (peri-

Gondwanan terranes). Paleomagnetic data places this Belt along the northern active margin of

Gondwana, most likely near or adjacent to the Amazonian Craton between c. 640-570 Ma (Fig.

2.1) (Samson et al., 2005; Balintoni et al., 2010 and references therein). At this time, the

Avalon Terrane consisted of one or more volcanic arc(s) and several arc-related basins and now occurs as a distinctive tectono-stratigraphic belt within the North Appalachians extending discontinuously from offshore eastern Newfoundland in Canada to southeastern New England in the USA (Nance, 1990; Thompson et al., 1996; Murphy et al., 1999; Thompson and

Bowring 2000; Murphy et al., 2004). The Boston Basin itself extends from northern Boston to the southern shore of Rhode Island and is bounded on the north, west and south by high-angle thrust faults; the nature of its eastern coastal boundary is unknown (Smith and Socci, 1990)

(Figs. 2.2 A, B).

Basement rocks that underlie the Boston Bay Group consist of the c. 625-600 Ma

Dedham and Westwood calc-alkaline granitoids (Hepburn et al., 1993; Thompson et al., 1996;

Thompson and Bowring, 2000), the younger Mattapan-Lynn Volcanic Complex (c. 597-596

Ma; Durfee Cardoza, 1987; Khoo et al., 2008) consisting of rhyolitic, andesitic, and basaltic flows, coarse granitic and syenitic dikes, volcanic breccias, scoria and lahars (Rast and Skehan,

1983; Socci and Smith, 1987; Thompson et al., 1996), and the mafic to intermediate flows, pillow lavas, and pyroclastics of the Brighton Volcanics (c. <587 Ma; Thompson and Grunow,

2004). The Mattapan-Lynn volcanic rocks record the oblique subduction of the Avalon Terrane beneath the northern Gondwanan margin and the development of magmatic arcs and associated back-arc basins within the Avalon Terrane between c. 640-570 Ma (Thompson, 1985; O‘Brien et al., 1996; Murphy et al., 1999). The Mattapan-Lynn Complex is interpreted as representing the vestigial remains of a caldera collapse (Thompson, 1985). The Brighton Volcanic rocks, also derived from the Avalon volcanoes, record a later phase of volcanic activity where subduction and ‗within-plate‘ continental rifting occurred simultaneously, typical of back-arc

19 basins (Cardoza et al., 1990). Based on the calc-alkaline nature of the basement granitoid bodies and the bimodal terrestrial volcanics and volcaniclastics of the Mattapan-Lynn Volcanic

Complex, the Boston Basin is interpreted as an arc-type basin (intra-arc or back-arc basin6) that developed between c. 640- 570 Ma (Skehan & Murray 1980; Hon and Hepburn, 1986; Nance,

1990; Socci and Smith, 1990; Cardoza et al. 1990).

Modern investigations of the sedimentary fill of the Boston Bay Group are constrained by a long standing formal stratigraphy based on ‗like-for-like‘ correlation from one outcrop to another and an antiquated assumption of a ‗layer cake‘ basin fill (Sayles and LaForge, 1910;

Sayles, 1914; Emerson, 1917; LaForge, 1932; Billings et al., 1939). Traditionally, the stratigraphy of the Boston Bay Group is described as two formal stratigraphic assemblages consisting of: (1) the Roxbury Conglomerate (c. 597 Ma: Ault et al., 2004) composed of a 150-

1300 m thick lowermost Brookline Member (dominantly conglomerate), a 180-500 m thick middle Dorchester Member (slate turbidites with minor conglomerate) and the 18-215 m thick upper Squantum Member (Bailey et al., 1976) of mostly diamictites with minor sandstones and siltstones, and (2) the overlying Cambridge Formation or ‗Cambridge Argillite‘ of Emerson

(1917) (c. 570 Ma; Thompson & Bowring, 2000) composed of 5.5 km of tuffaceous, finely- laminated argillite turbidites (Thompson and Bowring, 2000; Bailey and Bland, 2000). Each stratigraphic unit has been formally defined and recognized on the basis of the relative percentage of conglomerate, diamictite, sandstone and siltstone. However, even earlier workers noted that this scheme was simplistic (Billings et al., 1939; Dott, 1961; Bailey et al., 1976;

Socci and Smith, 1987).

6 See Appendix. 20

2.3 PREVIOUS INVESTIGATIONS AND AGE OF THE SQUANTUM MEMBER

Rocks belonging to the varyingly named Squantum Member (Billings, 1976, 1979),

Squantum Diamictite (Bailey, 1987) or Squantum ‗Tillite‘ (Sayles, 1914), were first described as ‗puddingstones‘, ‗conglomerates‘ and ‗slate conglomerates‘ by Crosby (1880). By comparison with then better-known Pleistocene facies, these rocks were then viewed as evidence of a glacial origin for the succession (tillites, glaciofluvial outwash and glaciolacustrine varvites respectively) by Sayles and LaForge (1910) and Sayles (1914, 1919).

A glacial interpretation of the Squantum Member was based essentially on the till-like appearance of the diamictites (e.g., poor sorting and lack of stratification) (Sayles, 1914;

Emerson, 1917; Billings, 1929; Billings et al., 1939). Rare, faintly striated pebbles were reported by Sayles (1914) but have never been verified by later workers (Dott, 1961; Billings,

1976; Bailey et al., 1984). In keeping with the glacial interpretation of the Squantum Member, lonestones in massive sandstones in the Newton area and laminated argillites at Squantum

Head (Fig. 2.2B) were identified as ‗ice-rafted dropstones‘ by Lahee (1914), Sayles (1914) and

Cameron (1979). Sayles (1919) and Caldwell (1964) interpreted the very thick (5.5 km)

‗Cambridge Argillite‘ as recording post-glacial glacioeustatic sea level rise and the development of a deep glacial lake. At that time, the Squantum Member was widely regarded as being Late Paleozoic in age and was consequently considered to be equivalent to the widespread Gondwanan glacial deposits of the southern hemisphere (e.g., ‗Dwyka tillite‘;

Sayles, 1914; Emerson, 1917; Billings, 1929; Billings et al., 1939). Wegener (1929) strongly disagreed with this model because the deposit lay well outside his reconstruction of Late

Paleozoic ice covers and he was later shown to be correct when the Squantum Member was dated radiometrically as Neoproterozoic in age (Lenk et al., 1982).

The Squantum Member was first described using a modern facies based approach by

Dott (1961) who identified a non-glacial setting for the Boston Bay Group. In his study he revealed that the clast lithologies of the diamictites and conglomerates were the same and that the Canbridge argillites were not glacial varvites but rather they were turbidites. As a result,

Dott (1961) interpreted the Squantum diamictites as mass flow deposits produced by the mixing of gravel and mud during downslope transport. This study was the catalyst for others that identified a more complex marine origin for many pre-Pleistocene ‗tillites‘ and the importance of debris flow processes and sediment mixing in creating till-like sediments within deep marine non-glacial successions (‗mixtites7‘; e.g., Crowell, 1964; Schermerhorn, 1966,

1974). Nonetheless, a glaciomarine interpretation of the Squantum Member still persists (Socci and Smith, 1987, 1990; Kawai et al., 2008; Hoffman and Li, 2009; Passchier and Erukanure,

2010) (see Carto and Eyles, 2011 for a detailed historical review of the Squantum Member).

A maximum age of 597 Ma has been determined for the Roxbury Conglomerate assemblage (weighted mean 207PB/238 U age based on four single grain analyses) from a volcanic ash associated with basalt interbedded with conglomerates at Rocky Neck, Hingham,

MA (Ault et al., 2004). Recently U-Pb data from a sandstone horizon associated with the diamictites revealed that the Squantum Member is younger than 593 Ma. The youngest age of the Roxbury Conglomerate assemblage at c. 570 Ma (207PB/206Pb date) is constrained by the youngest detrital zircon component from an ash bed in the overlying Cambridge ‗Argillite‘

7 Umbrella term used to describe rocks with a poorly sorted megaclastic lithology composed of a wide range of sediment size grades (diamictite) without regard to origin (Schermerhorn, 1966). 22

(Thompson & Bowring, 2000). Based on these considerations, the age of the Squantum

Member is widely regarded to lie somewhere between 593-570 Ma (Thompson and Bowring,

2000; Thompson et al., 2007). This confirms contemporaneous clastic deposition with

Avalonian arc volcanism of the Lynn-Mattapan Volcanic Complex (602-593 Ma; Ault,

2003).Given its Neoproterozoic age constraint, the Squantum Member is now considered by

‗Snowball Earth‘ proponents to be a key piece of evidence of a putative global or near-global correlative glaciation event (the ‗Gaskiers glaciation‘ at c. 580 Ma; Thompson et al., 2003;

Narbonne and Gehling, 2003; Thompson and Bowring, 2000; Bingen et al. 2005; Halverson et al., 2005; Kawai et al., 2008; Hoffman and Li, 2009).

2.4 SCOPE OF STUDY AND METHODS

The purpose here is to present a fully integrated sedimentological and basinal study of the Squantum Member and associated lithofacies in the context of the tectonic and volcanic setting of the Boston Basin. The fundamental basis of this reassessment consists of detailed lithofacies descriptions of more than 70 measured sections of the Squantum Member and associated facies exposed in outcrops throughout the Boston Basin (Fig. 2.2B). These are principally exposed in headlands along the Boston coastline and at several inland areas.

Individually, these outcrops are not very thick or laterally extensive but when considered together they provide a clear picture of the sedimentology of the Squantum Member and its associated strata and their origin and basinal context. Each outcrop was described (‗logged‘) using well proven field techniques (Miall, 1978; Eyles et al., 1983; Nicholas, 1999). A number of facies can be recognized on the basis of the variation in clast shape and abundance, together

23 with sedimentary structures such as the presence or absence of normal grading and stratification and syn- and post-depositional deformation structures. Special emphasis was placed on bed contacts and the vertical and lateral relationship between various facies.

Additional clast lithology and roundness data were collected (based on Pettijohn, 1957 values) and samples of rock removed for petrographic examination of thin sections and facies sub- types.

The most detailed investigation of the Squantum Member was conducted at Squantum

Head (Figs. 2.2A, B) which is the best exposure and the classic locality on which all previous interpretations of the Squantum Member, both glacial and non-glacial, are based (Sayles, 1914,

1919; Emerson, 1917; LaForge, 1932; Crowell, 1957; Dott, 1961; Lindsay et al., 1970; Bailey,

1984, 1987; Socci and Smith, 1987; Passchier and Erukanure, 2010). Other additional information was obtained from deep drilling projects in the Greater Boston urban area (Billings and Tierney, 1964; Billings, 1976, 1979).

This study concludes by discussing the global significance of the results for reconstructing the Neoproterozoic paleoclimate and paleoenvironments and testing the validity of the SEH.

2.5 DESCRIPTION AND INTERPRETATION OF LITHOFACIES

Strata of the Boston Bay Group consist of five principal lithofacies found in conformable successions and are thus genetically related. These are: (1) conglomerate, (2) sandstone, (3) diamictite, (4) argillite and (5) volcanic-sedimentary facies. Each is described and interpreted in turn below.

2.5.1 Conglomerate facies

Conglomerates constitute a considerable cumulative thickness portion of the entire basin fill, approximately 1.3 km out of 7 km (Billings and Tierney, 1964; Tierney et al., 1968).

These facies occur as massive and crudely stratified clast-and matrix-supported units interbedded with massive, graded and cross-laminated sandstones and rarely diamictites.

Massive, clast-supported conglomerates

Massive conglomerates are typical of rocks traditionally labeled as the ‗Brookline

Member‘ (Billings, 1976, 1979; Socci and Smith, 1987) occurring as single or multi-storey stacked beds with erosive and scoured bases and irregular upper boundaries. These facies are internally structureless (massive) and moderately to well sorted and occur as thick (up to 5 m thick) beds of tightly packed clasts with a low percentage of brown-grey coloured coarse silty- sand matrix (< 10 volume % of rock) (Figs. 2.3A, 2.8E). Conglomerates are typically composed of pebble-to boulder-sized clasts (4-40 cm in diameter). Clast shape ranges from subrounded to well rounded and clasts are dominated by basement-derived volcanic lithologies such as rhyolite, andesite, basalt, diorite, and some granite, although the relative abundance of each clast type is highly variable between the outcrops. The long axes of clasts are parallel or subparallel to bedding planes in some outcrops. Thin-section analysis revealed that the matrix is composed of medium- to coarse-grained subrounded to rounded grains of quartz with lithic clasts up to 5.0 mm in diameter (75 vol. % of sample) and volcanic fragments dominated by red, gray, and white felsites and dacite (25 vol. % of sample). It is common for conglomerate

25 beds to be found in conformable association with thick beds (up to a metre-thick) of massive and thinly laminated, normally graded fine-to very-fine sandstone beds (described and interpreted in section 2.5.2) having exactly the same composition as the matrix material of enclosing conglomerates. The bases of the sandstone beds are gradational to underlying conglomerate. Some ‗disrupted‘ conglomerate beds are composed of granule-to cobble-sized clasts (< 5 cm in diameter) and contain deformed rafts of laminated argillite facies (Figs.

2.3B,C and 2.8D). Individual beds are laterally continuous for the extent of exposures (for several metres).

Massive, matrix-supported conglomerate

These facies have the same clast lithology and average bed thicknesses (Figs. 2.3D, E,

2.8F) as clast-supported facies with which they are commonly interbedded but clasts are less well rounded (subangular to subrounded) and are supported by a larger amount of brown-grey sandstone matrix (25-30 vol. % of rock). The average clast size is generally smaller and more uniform in size (3-12 cm in diameter) compared to clast-supported facies. In thin section, the conglomerate matrix consists of a cryptocrystalline, sericitic (fine-grained mica either as muscovite or illite) groundmass with abundant very fine-to-fine quartz grains (70 vol. % of matrix), with secondary (25 vol. % of sample) volcanic grains and small (< 1 cm) subrounded porphyritic rhyolite and diorite clasts (5 vol. % of sample). Matrix-supported facies are generally massive but in some cases display a crudely developed grading at the base of beds.

Some units contain contorted lenses and rafts of coarse sandstone (Figs. 2.3E, 2.8F). Individual beds are laterally continuous for the extent of exposures (for tens of metres).

Matrix-supported conglomerates are interbedded and overlain by lenticular and tabular shaped beds up to 2 m thick of coarse-grained feldspathic to lithic (volcanic, plutonic sedimentary and metasedimentary) to quartz-rich sandstone having exactly the same composition as the matrix material of enclosing conglomerates. The sandstone beds are commonly massive but in some cases exhibit low-angle cross-lamination. The bases of the sandstone beds are gradational to underlying conglomerate.

Horizontally- stratified conglomerates

Well-to-crudely-stratified conglomerates are defined by the repetition of 30-60 centimeter-thick horizontal beds of gravel and pebble conglomerate with beds of coarse- grained granule-sandstone that range in thickness from 5-75 cm (Figs. 2.3F, 2.8G). Clast and matrix lithologies are similar to other conglomerates but clast size is reduced (1-5 cm in diameter). The long axes of clasts are parallel or subparallel to bedding planes. Overall, the clast size appears to decrease towards the top of a bed as successive conglomerate layers become finer grained and thinner, creating ill-defined stringers of clasts and a crude stratification. Stratification appears to vary from plane parallel to sub-parallel. This facies occurs as thick sheet-like bodies typically within multistorey sequences that range from 1 to 10 m in thickness; beds blend into each other without sharp boundaries. Individual units are seldomly laterally continuous, extending 5 to 10 metres.

Interpretation

The characteristics of massive and clast-supported conglomerates are entirely consistent with deposition from sediment-gravity flows on a subaqueous slope. These facies are typical of cohesionless gravelly debris flows wherein momentum transfer and clast buoyancy was maintained by mutual interaction of clasts or limited matrix buoyancy (e.g., Walker, 1975;

Nemec and Steel, 1984; Sohn et al., 1997; Sohn, 2000a, 2000b; Lowe and Guy, 2000; Mulder and Alexander, 2001; Sohn et al., 2002; Tripsanas 2008). Low proportion of matrix and erosional and scoured bases support this contention (Mulder and Alexander, 2001). The alignment of clasts oriented parallel and subparallel to bedding plane is also a feature indicative of laminar flow (Prior, 1990; Mohrig et al., 1999; Eyles and Eyles, 2000;

Dowdeswell 2001; Mulder and Alexander 2001).

The best-known mechanism for the deposition of the laminated and normally graded and massive sandstone beds found in conformable association with clast-supported conglomerates is by means of turbidity currents or concentrated-density flow deposits (Mulder and Alexander, 2001). In some cases, similar sandstone beds may be found with subaerial debris flows resulting from turbulent fluidal flow or heavily sediment-laden stream flow following the debris, but in these cases its contact with the underlying debris flow will be sharp

(Nemec and Steel. 1984). The gradational contacts between sandstones and underlying conglomerates and the similarity in the petrographic composition of the sandstone beds and the matrix and clasts in the associated conglomerate facies suggest that these sand-laden turbulent flows were entrained above and deposited after the moving debris mass. The preservation of graded bedding in the sandstone beds implies that the water was sufficiently deep to prevent

28 currents induced by wind, waves and tides from reworking the newly deposited sediment

(Hampton, 1972; Nemec et al., 1984; Piper et al., 1992; Mohrig et al., 1999; Mulder and

Alexander, 2001). Clasts within clast-supported conglomerates are relatively well sorted given the narrow range of clast sizes and clast lithologies indicate intrabasinal sources from local felsic and granitic basement rocks. The predominantly rounded nature of the clasts clearly suggests an original fluvial source for the sediment prior to redeposition downslope.

Matrix-supported facies are typical of cohesionless subaqueous debris flows able to incorporate matrix by mixing of pebbles and cobbles with large volumes of sand during downslope flow (e.g., Mulder and Alexander, 2001; Piper et al., 2004; Deynoux et al., 2005;

Frey-Martinez et al., 2006; Tripsanas et al., 2008). In these flows clasts were supported by the frictional strength of the matrix as well as clast collisions (Sohn et al., 1999; Mulder and

Alexander, 2001). Tectonically unstable areas are conducive to sediment mixing and local redeposition, resulting in matrix-rich conglomerate beds (Larsen and Steel, 1978). Crude grading of the matrix-supported conglomerates indicates the onset of fluid turbulence which typically evolves downslope in resedimented conglomerates (e.g., Walker, 1975; Nemec and

Steel, 1984; Trop et al., 1999). As observed in clast-supported conglomerate facies, the massive sandstone beds found in conformable association with matrix-supported conglomerates have the same petrographic composition as the matrix in the associated conglomerate facies and therefore it is likely they represent high-concentration turbidity currents (e.g., Bouma, 1962; Eyles et al., 1993; Trop et al., 1999; Kneller and Buckee, 2000) or concentrated-density flow deposits (Mulder and Alexander, 2001). These sandstone facies are discussed in further detail below (section 2.5.2).

Better-sorted, horizontally-stratified conglomerates are typical of more distal submarine debris flow deposits and/or high-density turbidites having a traction carpet phase of deposition (Walker, 1975; Miall, 1985; Hiscott, 1994; Sohn et al., 1997). A sub-parallel to parallel alignment of clasts to the bedding plane is a feature typical of coarse-grained sediment deposited from traction at the base of turbulent flow (Lowe, 1982). It is notable that stratification is generally very crude, and bed boundaries appear to be amalgamated, which suggests fluctuating or pulsating deposition from fluidized flow (Harms et al., 1975).

‗Disrupted‘ conglomerate facies are characteristic of poorly mixed slump and/or slide deposits on the more proximal parts of unstable subaqueous slopes (Walker, 1975, 1984; Sohn,

2000b; Mulder and Alexander, 2001; Sohn et al., 2002). Disruption occurs when slumps entrain underlying bed materials and become fluidized by incorporating water from the overlying water column eventually causing the bed to lose any original bedding or coherence downslope (Hampton et al., 1996). Differential loading of masses of gravel into water- saturated sand may also cause deformation and trigger mixing (Chough et al., 1990). These facies likely represent proximal upslope facies compared to the better-sorted and more homogeneous matrix- and clast-supported types (see Walker, 1975; 1984; Sohn, 2000a, 2000b;

Mulder and Alexander, 2001; Sohn et al., 2002).

Overall, the characteristics of the conglomerate units indicate that they were emplaced on the submarine slopes of the basin at intermediate to deep water depths as evidenced by their association with turbidites and the presence of crudely stratified, graded and slumped and disrupted conglomerate facies (Walker, 1975; 1984; Nemec and Steel, 1984; Sohn, 2000a,

2000b; Mulder and Alexander, 2001; Sohn et al., 2002).

2.5.2 Sandstone facies

Sandstones occur as massive, normally graded and cross-laminated coarse- to fine- grained facies as well as deformed facies. Overall, petrographic composition of the sandstone units is the same as the conglomerates, including subrounded-to-rounded feldspathic, lithic

(volcanic, plutonic, and sedimentary) and quartz-rich grains.

Massive sandstone

Massive thinly- (0.7 to 1 m thick) and thickly-bedded (up to 9 m thick) medium-to fine- grained sandstones are very common. Individual beds are laterally continuous for the extent of exposures (for tens of metres). Bed boundaries are relatively flat or gradational. Sedimentary structures are absent except for stringers of coarse-grained sandstone (Fig. 2.4A). Very rare isolated pebble-sized stones (< 3 cm in diameter) are present in some beds. Some massive sandstone units transition into sandstones that display crudely defined, low-angle cross- lamination (Fig. 2.4B). The units occur as tabular bodies (up to 0.5 m thick) and are always found interbedded with matrix-supported conglomerates, normally graded sandstones and massive diamictites (Fig. 2.8B).

Laminated and normally graded sandstone

These units occur conformably interbedded with conglomerate facies and consists of thick successions of graded, planar laminae (> 8 mm in thickness) of fine-grained sandstone

31 which grade upwards into siltstone. The composite thickness of laminated sandstone facies ranges from one metre to tens of metres (Figs. 2.4C/ 2.8I). Individual laminae frequently show loaded bases; wavy and convolute laminae are common. Laminae are laterally persistent in outcrop (several tens of metres). Soft sediment deformation features such as large slump features, overturned folds (amplitudes 2 cm to 2 m) and pinch-and-swell bedding are typical

(Figs. 2.4D, E ). Massive and graded sandstone beds are commonly disrupted by concave- upward lenses (‗dishes‘) up to 30 cm wide, of medium to coarse-grained sandstone and subvertical columns (‗pillars‘) up to 4 m in height (Fig. 2.4F).

Interpretation

As with conglomerates, sandstones record the downslope transport of sand by turbulent flows. Massive sandstone facies are typical of the lowermost ‗Ta‘ division of turbidites recording transport by high-concentration turbidity currents and the rapid en masse ‗freezing‘ sufficient to suppress tractional transport (e.g., Bouma, 1962; Eyles et al., 1993; Trop et al.,

1999; Kneller and Buckee, 2000) or concentrated-density flow deposits (Mulder and

Alexander, 2001). Rare lonestones within massive sandstones are typical of clasts transported along traction carpets by high-concentration turbidity currents (Postma et al., 1988; Stow and

Johansson, 2000). Overlying low-angle cross-laminated sandstones are typical of ‗Tb‘ and ‗Tc‘ divisions of Bouma (1962), representing high-density turbidity current deposits settled from traction currents (Lowe, 1982; Ineson, 1989; Socci and Smith, 1990 Walker and Plint, 1992).

The thick successions of laminated, normally graded sandstones with convolute and wavy laminations are also typical of classic ‗Ta‘ turbidity currents (Walker, 1992; Eyles et al.,

1993; Trop et al., 1999; Kneller and Buckee, 2000; Mulder and Alexander, 2001). The absence of upper 'Bouma' divisions of turbidites (‗Td‘ and Te‘ divisions) most likely reflects their proximity to the source, an apparent discontinuity of the flow regime downcurrent, or post- depositional erosion by succeeding turbidity currents (Bass, 2004). ‗Dish-and-pillar‘ structures are typical of well-known water-escape features produced by the rapid deposition of sand by mass flow resulting in abrupt upward release of pore waters as sandstone dikes (e.g., Lowe and

LoPiccolo, 1974; Lowe, 1982; Rossetti, 1999).

2.5.3 Diamictite facies

Diamictite facies (‗Squantum Tillite‘; Sayles, 1914) are represented by ‗homogeneous‘

(massive) and ‗heterogeneous‘ (chaotic) diamictite beds. Significantly, both facies types are conformably interbedded with laminated and normally graded argillites, massive conglomerates and sandstones.

Homogeneous massive diamictite

The diamictites of this facies type show dispersed clasts supported by a poorly sorted matrix of gritty sandy mudstone, accounting for as much as 70-80 % of the total rock volume (Figs.

2.5A, 2.8A-C). Diamictite facies are remarkably similar to the matrix-supported conglomerate facies described in section 2.5.1 in both outcrop appearance and in sedimentology; however, the two facies types are differentiated in this study based on the composition and abundance of the matrix. Matrix-supported conglomerates are composed of densely concentrated clasts embedded in a coarse-to medium-grained sandy matrix (60-70 vol. % of rock), whereas the

33 matrix of the diamictite units is composed of a greater abundance of clay and silt (mud), and a smaller volume of clasts (20 to 30 % vol. of rock).

Clasts ranging in size from 8 to 25 cm in diameter are mostly subrounded to subangular, with the smaller size fractions (< 6 cm in diameter) typically being better rounded.

Clast size and shape distributions are similar to that observed in conglomerates; clasts up to 80 cm in diameter are present but rare. Clast lithologies are exclusively intrabasinal including multi-coloured felsic and mafic volcanics (up to 45 vol. % of rock) with lesser granodiorite, granite and quartzite clasts (35 vol. % of rock). Intraclasts of massive, graded sandstone and laminated argillite up to 0.75 mm in diameter are common (20 vol. % of rock). Measurement of clast orientation in lithified rock is not straightforward but Lindsay et al. (1970) reported a clear trend (N68 o- N73 o) in long axis orientation of one hundred clasts from diamictite beds at

Squantum Head. No striated, faceted, or bullet-shaped clasts were found within any outcrop of the diamictite, despite observation of many several metres of thickness and areas of broken rubble where numerous clasts were fully exposed and completely freed from their matrix.

Workers who earlier identified ‗striated‘ clasts admitted that striae are faint and rare (Sayles,

1914); these are possibly the product of the abrasion of clasts by the sandy matrix (see

Schermerhorn, 1974).

In thin section, the dense, cryptocrystalline matrix of the diamictite units (83 vol. % of sample) contains large numbers of irregularly scattered, subrounded lithic clasts up to 4 mm in diameter, composed of crystalline quartz (2 vol. % of sample), intermediate volcanic clasts (5 vol. % of sample), granitic rocks clasts (3 vol. % of sample), volcanic glass (1 vol. % of

34 sample), and subordinate sericitised feldspar and very small grains of iron oxide (6 vol. % of sample). Thin stringers of fine sand up to 3 mm thick commonly traverse the diamictite matrix.

In outcrop, diamictites occur as planar tabular beds up to 8 m thick, although they are reported as being up to 215 m thick in the subsurface (Billings and Tierney, 1964; Tierney et al., 1968). At the type section at Squantum Head, diamictites are conformably interbedded with parallel-laminated, normally graded argillites and laminated pebbly argillites (the ‗glacial varves‘ of Sayles, 1919) (refer to Figs. 2.6A-C, 2.8C) and massive muddy sandstones. Rarely, a crudely developed ‗coarse-tail‘ normal grading is shown by massive diamictite beds with higher concentrations of clasts toward their basal parts (Fig. 2.5B). In general, upper bed boundaries are mainly gradational/ irregular to sandstones and conglomerates, whereas lower bed boundaries are often gradational or disrupted by loading.

Heterogeneous diamictites

A significant proportion of the diamictite units measured in outcrop are of the heterogeneous or ‗chaotic‘ type. These have exactly the same matrix type, clast lithology, clast orientation and clast size/shape characteristics as homogeneous facies but have a crudely mixed and chaotic internal structure consisting of wisps and stringers of sandstone and deformed and ‗chaotically-mixed‘ rafts of clast-supported conglomerate and massive muddy sandstone matrix (Figs. 2.5C, D and 2.8C). Rafts of deformed laminated argillite facies and rounded mud slabs are common (Figs. 2.5E). Slump structures are also common.

Heterogeneous diamictite facies occur as tabular beds (1 to 7 m thick) or as lenticular, thin beds ranging from 10 to 40 cm in thickness. Both heterogeneous and homogeneous diamictite

35 units are interbedded at Squantum Head. Gradational/ irregular upper bed boundaries and loaded bases are typical.

Interpretation

The most important outcrop observation is that the diamictite facies differ from conglomerates simply by having a greater proportion of matrix (see also Dott, 1961, p. 1300) indeed; the diamictites were described as ‗polymictic conglomerate‘ by Bailey et al., (1976) and Socci and Smith (1990). The clast size, lithology and shape characteristics are identical to the conglomerate facies indicating a common source. Consequently, diamictites are interpreted as being genetically related to conglomerates, produced by the mixing and subsequent ‗textural inversion‘ of gravel and muddy matrix material during dowslope flow. This is the principal means by which diamictites are produced regardless of sedimentary environment (see Crowell,

1957; Dott, 1961; Folk, 1974; Nemec and Steel, 1984; Nemec et al., 1984; Eyles, 1990; Eyles and Eyles, 2000; Mulder and Alexander, 2001; Mohrig & Marr, 2003). Schermerhorn (1974) referred to such deposits as ‗mixtites‘. Such facies are known to develop when primary fan or slope conglomerates are reworked by gravity and subsequently deposited on and mixed with soft and water-saturated fine-grained sediment in the middle/lower reaches of the fan/slope

(Crowell, 1957; Folk, 1974; Walker, 1975, 1984; Sohn, 2000b; Mulder and Alexander, 2001;

Sohn et al., 2002). As described in Crowell‘s (1957) model for the formation of the pebbly mudstone deposits of Californian mélanges, if the bulk densities of such deposits are great enough in relation to steepness of the slope, the mud and gravel mixture will continue to move downslope basinward, resulting in the churning, mixing and dispersing of pebbles throughout

36 the mud. If the mixed mass is viscous it may plough into more substratum causing slabs of underlying mud to be rolled up and incorporated in the mass to form slump overfolds and irregular masses and lenses. Although the presence of appreciable mud in the matrix of the diamictites would have led to an eventual increase in the viscosity of the flows, the presence of stratified and normally graded diamictites beds indicates that most of the flows remained fluid

(Mulder and Alexander 2001). Moreover, as Crowell (1957) emphasized, mud with a high content of clay may transform from its gelled and solid state with motion due to the breakdown of the thixotropic barrier. As a result, the gelled structure with the mud breaks down

(electrostatic bonds between clay particles are broken) causing the material to become fluid.

The sudden deposition of gravel may also cause a breakthrough of the thixotropic barrier so that it moves and slides easily (Boswell 1948; Montenat et al., 2007). As the flow decelerates and eventually attenuates, the mixture transforms thixotropically, so that the flows become viscous and ‗pasty‘ inhibiting further mixing of the debris mass, such that slumped and soft- sediment features are preserved within the rock and isolated masses of the original sediment components (mud and coarse clastics) can still be recognized in outcrop. Furthermore, it has been shown that both clay-rich and clay-poor subaqueous debris flows undergo marked stretching due to the intrusion of ambient water, leading to the eventual break-up and disintegration of the debris flow into smaller pieces (Mohrig et al., 1998; Elverhøi et al., 2002;

Ilstad et al., 2004a, b, c). If the flow is moving slowly, the friction remains relatively high and the flow attenuates before most of the mass can be completely remolded (De Blasio et al.,

2004) again, isolated masses of the original sediment components will be observable in outcrop.

The applicability of Crowell‘s model to the Squantum Member is confirmed by the great variability in texture and varying degree of mixing observed in the outcrops of the

Squantum Member (i.e. sections of undigested and unmixed conglomerate and mud, slump overfolds and load casts). Heterogeneous diamictite facies showing masses of conglomeratic material and argillites ‗capture‘ the initial stages of downslope mixing and provide a ‗missing link‘ between clast-supported conglomerate facies and the more highly evolved homogeneous diamictite facies produced by downslope mixing. They are typical of slumped deposits that result from the downslope sliding and slumping of heterolithic sediment along highly unstable basin margins (Eyles et al., 1993; Tripsanas et al., 2008; Akdağ and Kirmacia, 2008).

Interbedding of diamictites with thinly laminated, normally graded argillite beds, interpreted below as turbidites (section 2.5.4) provides key contextual evidence of a subaqueous depositional environment.

2.5.4 Argillite facies

The term ‗argillite‘ has been applied to rhythmically laminated (0.1 to 1 mm thick) muddy-siltstone facies that grade subtly into mudstone (Figs. 2.5F, 2.8H). This is the dominant facies type in the northern portion of the Boston Basin where they reach total thicknesses of as much as several kilometers and is formally described as a separate stratigraphic unit (the

‗Cambridge Argillite‘); in reality, such facies are interbedded with all other facies throughout the basin typically having thicknesses of several centimeters to tens of metres.

Laminations are parallel to wavy to convolute and are laterally continuous (several tens of metres) in outcrop. Contacts between individual lamina are normally sharply defined.

Starved ripples and subtle small-scale cross lamination are common. Soft sediment deformation structures, such as small-and large-scale slump folds, are ubiquitous often occurring within discrete horizons within thicker successions. At the type section of Squantum

Head, these facies are intimately interbedded with diamictites and also occur as a separate thick succession (up to 6 m thick) containing thin (5-12 cm) beds of coarse-grained volcanic lapilli tuff (see section 2.5.5 below and Figs. 2.7A, B, 2.8H). Petrographic analysis reported by

Thompson and Bowring (2000) and those reported here, show that laminated argillites are composed of quartz, plagioclase, feldspar, melaphyre fragments, and epidote, all indicative of a local volcanic/igneous source area; with apatite, monazite, rutile and ziron as accessory minerals (80 vol.% of sample). In addition, some 20 % of the volume is composed of reworked volcanic ash consisting of very fine elongated, dark brown coloured grains of unweathered volcanic glass indicating a very local volcanic source area (Fig. 2.7D). It is noted that glass is metastable and readily alters to clays and zeolites, and therefore is not readily preserved (Fisher and Schmincke, 1992).

Laminated pebbly argillites within diamictites at Squantum Head are very thin (< 2 cm) and contain numerous small clasts of a restricted diameter (< 3 cm) (Figs. 2.6A-C). These facies have previously been interpreted as ice rafted ‗dropstones ‘within varvities (see below).

Qualitative observation shows that larger clasts tend to be subrounded whereas smaller clasts are predominantly well rounded. Significantly, clasts are of the same lithologic composition as those present in associated diamictites and conglomerates (namely crystalline quartz, granitic rocks, felsic volcanics, and fragments of siltstone and quartz). The matrix is also the same cryptocrystalline, sericitic-rich matrix present in diamictites. Observations of cut slabs and thin-sections show that clasts preferentially occur toward the base of individual layers defining

39 a crude grading, or ‗float‘ within the layers (see Chapter 3 of this thesis for a detailed description). Depression of laminae below clasts is rare and clasts seldom penetrate into underlying laminae.

Interpretation

Laminated and normally graded argillite facies are turbidites as well established by

Dott (1961), Thompson and Bowring (2000) and Bailey and Bland (2000). The primary signature of turbidity current deposition is the presence of graded bedding (e.g., Prior and

Bornhold, 1990; Eyles and Eyles, 2000; Lowe and Guy, 2000; Mulder and Alexander, 2001).

A relatively deep marine ‗below wave base‘ origin is indicated by their paleobiology.

Clustered and solitary Ediacaran microfossils (Bavlinella sp.) occur in argillite facies at

Hewitt‘s Cove in Hingham, MA (Fig. 2.2B) (Vidal, 1976; Lenk et al., 1982; Thompson and

Bowring, 2000). In addition, in situ specimens of Ediacaran fossils, interpreted as possibly

Aspidella terranoviva (ring fossils) occur elsewhere within silt-rich turbidites (Bailey and

Bland, 2000; Bailey, 2005). Passchier and Erukanure (2010) report two dislocated stromatolite hemispheroid specimens at Squantum Head possibly reworked downslope into deep water by turbidity currents. The presence of soft sediment structures at discrete intervals within argillites indicates that the depositional environment was subject to episodic sediment instability typical of seismically active settings (Bowman et al., 2004 and references therein).

Thin layers of pebbly argillite within laminated, normally graded argillites at Squantum

Head were interpreted as a record of ice rafting (varvities) by Sayles (1914, 1919), Sayles &

LaForge (1910) and Cameron and Jeanne (1976) and Socci and Smith (1987) and Smith and

Socci (1990). Large isolated clasts regarded as possible ‗dropstones‘ were also identified within sandstone and conglomerate units by Cameron and Jeanne (1976) but have been reinterpreted as being emplaced by turbidites (Bailey, 1984), as is well-known elsewhere (e.g.,

Nemec and Steel, 1984; Postma et al., 1988; Crowell, 1999; Benvenuti, 2003; Eyles and

Januszczak, 2004). Laminated pebbly argillites are interpreted here as thin debrite horizons that escaped from thicker debris flow deposits (diamictite) upslope. They are directly comparable to the ―co-genetic‖ debrite-turbidite beds of Talling et al. (2004) documented in distal basin- plain settings where debris flows evolve into turbulent silty flows due to flow separation effects (e.g., Hampton, 1972; Barker et al., 2008; Pyles and Jennette, 2009)(see Chapter 3).

2.5.5 Volcanic-sedimentary facies

Petrographic data shows that argillites are composed of much reworked volcanic ash consistent with an arc basin setting (see above). Field observations and petrographic analyses also reveal the presence of distinct beds of reworked massive and crudely graded lapilli tuff within argillites at Squantum Head; previous work simply described these beds as thin

‗conglomerates‘ (e.g., Socci and Smith, 1987; Bailey et al., 1989; Passchier and Erukanure,

2010). Reworked tuffs are thin (between 5 and 12 cm thick) and have a lensate, channellized bed geometry with loaded and erosional bases (Figs. 2.7A, 2.8H). They are dominated by abundant lithic clasts of very fine-grained volcanic rock that range from 1-3 cm in diameter consisting of very fine felsic (rhyolite) clasts and intermediate volcanic clasts (dacitic) (40 vol.

% of rock). Also present are subrounded crystalline clasts of quartz (30 vol. % of sample), iron oxides (16 vol. % of sample), alkali feldspar (10 vol. % of sample) and plagioclase (4 vol. % of

41 sample) (Fig. 2.7C). Discrete ash beds of dacite to rhyodacite composition in graded argillite

(‗Cambridge Argillite‘) have been previously identified at the Mystic Quarry in Somerville

(Weisner et al., 1995; Thompson & Bowring, 2000) (Fig. 2.2B).

In general, volcanic rocks are known to constitute a significant portion of the fill of the

Boston Basin. For instance, the Brighton Volcanics (represented by mafic to intermediate flows, pillow lavas, and pyroclastics) are complexly interbedded with the clast-supported conglomerates (e.g., Kaye & Zartman, 1980; Zartman & Naylor, 1984; Thompson, 1993). The most informative outcrop exposure of this relationship occurs on the coast at World‘s End

Reservation Park (Fig. 2.2B) where massive conglomerate facies are interbedded with Brighton pyroclastic flows, which rest unconformably on the Dedham Granite (Fig. 2.9). Another exposure at Nantasket Beach (Fig. 2.2B) consists of granitic and volcanic basement rocks

(Dedham Granite, Lynn volcanics and older volcanics) overlain by Mattapan-basaltic lava flows and water-lain tuffs interbedded with massive, clast-supported conglomerate facies

(Thompson, 1993).

Beds of massive and crudely graded lapilli tuff are well documented from modern volcanic settings where lapilli-sized ash is reworked downslope subaqueously (e.g., Fisher,

1984; Ayres et al., 1991; Trop et al., 1999). Deposition is by non-cohesive sediment-gravity flows (Ayres et al., 1991; Mulder and Alexander 2001). This is evidenced by the presence of grading, low proportion of matrix, erosional bases and random orientation of clasts (Mulder and Alexander, 2001).

2.6 DEPOSITIONAL AND TECTONIC SETTING

Interfingering of clastic sedimentary and volcanic lithologies in the Boston Basin and the presence of subaqueously reworked tuff and argillitic turbidites of reworked ash provide key evidence for ongoing volcanism contemporaneous with subaqueous sedimentation in the

Boston Basin. Observations and interpretations made from outcrops confirm the findings of

Dott (1961) and, in most regards, those of Socci and Smith (1987) in terms of the importance of subaqueous sediment-gravity flow processes during deposition of the Squantum Member and the associated strata of the Boston Basin. Considered together, the sedimentology, stratigraphy and sequence context of the conglomerate, heterogeneous and homogeneous diamictite, sandstones and argillite facies suggests that these facies represent a continuum of genetically related sediment-gravity flow facies recording the downslope movement and mixing of gravel and mud (Fig. 2.10). Complete lithologic gradations between conglomerate and diamictite units can be observed in several outcrops (Fig. 2.11). As outlined above,

(section 2.5.3) this model agrees with Crowell‘s (1957) work on the origins of ‗pebbly mudstones‘. The term ‗mixtite‘ introduced by Schermerhorn (1974) to describe such facies is perfectly applicable to the Squantum diamictites.

Based on the results of this facies analysis it can be deduced that the paleoenvironment of the basin was one in which slumps, debris flows and turbidity currents were actively transporting large volumes of fluvially-transported and volcanically derived sediment into a deep marine and mud-rich slope environment (Fig. 2.12). A muddy depositional system is indicated by the considerable thickness of laminated, normally graded argillite turbidites (‗Cambridge Argillite‘) and muddy diamictites, most likely reflecting the

43 availability of large volumes of easily weathered volcanics and ash. The absence of features such as slump scars precludes an upper slope setting for these deposits. The presence of deep water Ediacara fauna in associated argillites confirms a deep water basin environment.

Evidence of loading, sliding and slumping of sediment in the form of folds, slumped features, soft sediment deformation, and dish and pillar structures are ubiquitous, possibly reflecting an active tectonic setting and the remobilization of sediment in response to seismicity.

Furthermore, the composition of the sedimentary rocks of the basin fill indicates a local source from the arc-related rocks of the Avalon Terrane (felsic rocks of the Lynn-Mattapan Complex, quartzites of the Westboro Formation and the Westwood and Dedham granitoids) (Dott, 1961;

Socci and Smith, 1987).

Regional mapping provides significant insights into the complex nature of sedimentation within this intra-arc (Nance, 1990; Socci and Smith, 1990) or a back-arc

(Cardoza et al., 1990) basin. The Brighton Volcanics and Roxbury-type conglomerates appear to have developed together in an initial series of rifts now filled by the stratigraphically lowermost Lynn-Mattapan Volcanic Complex and bounded by east-northeast-trending syn- volcanic/ syn-depositional normal faults such as the Mount Hope, Savin Hill, and Eliot faults

(Billings, 1929, 1976; LaForge, 1932; Zen, 1983; Thompson, 1993). The considerable thickness of the basin fill (~ 7 km thick) implies appreciable and rapid subsidence. The northern part of the basin is dominated by fine-grained facies of the ‗Cambridge Argillite‘

(Crosby 1880; Socci and Smith, 1987; Bailey, 1987; Thompson, 1993) whereas the thickest accumulations of conglomerates occur along the opposing southern margin. Paleocurrent data collected throughout the basin by Socci and Smith (1987), as well as slump-fold axes and pebble axis data collected by Lindsey et al. (1970) confirm a dominant northeastward direction

44 of sediment transport. It is significant that diamictites are found only near the basin‘s southern margin (e.g., in upslope positions proximal to sediment source areas receiving large volumes of fluvial conglomerates, gravels and fine grained sediment).

The bulk of the succession is dominated by laminated argillite facies (~5.5 km of a total basin fill thickness of ~ 7 km) recording prolonged turbidity current activity. Conglomerates, diamictites and sandstone facies units comprise the remaining ~1.5 km of the basin fill recording the episodic delivery of coarse debris into the basin by a variety of sediment-gravity flows including slumps, debris flow, and turbidity currents. The sedimentology and stratigraphy of the Boston Basin fill resembles that of modern and ancient back-arc basins where sources of sediment include pelagic fall-out, airborne ash, and submarine gravity flows derived from the adjacent volcanic arc. Sedimentation is such basins are strongly influenced by the steep, unstable nature of the volcanic arc flanks, the nature of the basin‘s boundary faults, and the coeval volcanism. The resulting basin fills consist of debris flows and silty basinal turbidites interbedded by pyroclastic and volcanic flows that often form complex submarine- fan or apron depositional systems (Karig and Moore, 1975; Busby and Ingersoll, 1995; Taira,

2001; Brandes et al., 2008).

The identification of diamictites at Squantum Head as tillites (‗Squantum Tillite‘) by

Sayles (1914) and Sayles and LaForge (1910) carried with it the implication that such facies recorded a single climatic event and formed a single stratigraphic unit across the basin. This reasoning in turn, underlay the concept that such facies could be used for precise lithostratigraphic correlation (Billings et al., 1939, p. 1873) within a simple ‗layer cake‘ basin fill. Billings et al. (1939. p. 1873) argued that ‗the tillite at the time of deposition was practically horizontal for the Boston Basin as a whole.‘ Dott (1961) and Socci and Smith

(1987) have strongly argued against the continued usage of the pre-existing formal stratigraphic subdivisions, though this conflicts with the very simplified layer cake stratigraphy presented in their depositional model (see Socci and Smith, 1987; p. 34). In reality, all the lithofacies described above occur repetitively at random intervals within the succession, including the diamictites.

The general trend towards a reduction in grain size from the southern to the northern margin of the Boston Basin combined with paleocurrent directions from clasts and slump folds suggests that sediment was shed downslope into the basin across relatively complex submarine fans by mass flow processes. Moreover, the basin fill is replete with lithofacies typical of a submarine fan regime (Miall, 1978; Pickering et al. 1986, 1989; Mutti 1992; Ross et al. 1994;

Pickering and Corregidor, 2005; Akdağ and Kırmac, 2008). However, given outcrop limitations it is not possible to identify any definitive evidence of channeling of facies suggestive of a classic submarine fan setting (see Richards et al., 1998). ‗Normal‘ fan growth may have been disrupted by the complex structure of the basin, rapid basin subsidence and the rapid influx of sediment from several sources that varied spatially and, most likely temporally.

Tectonic disruption to ‗normal‘ fan growth has been documented in both modern and ancient arc-related basins and trenches due to instability associated with rapid uplift, subsidence, and/or horizontal movement along fault zones within the basin (Stow et al., 1985; Carey and

Sigurdsson, 1984; Fisher and Schmincke, 1994; Richards et al., 1998).

Interfingering of the clastic sediments of the basin fill with the Brighton Volcanics provides key evidence that volcanoes were active, emitting ash and pyroclastic flows and magma during deposition. Significantly, in some areas the Brighton Volcanics are represented by stratified water-lain tuffs (Bell, 1948; Hon and Hepburn, 1986; Thompson, 1993) indicating

46 that the basin transitioned into a marine environment early in its formation. Compounding evidence of a volcanic influence on sedimentation is presented by the newly recognized reworked lapilli-tuff beds within laminated and normally graded argillites at Squantum Head.

2.7 DISCUSSION

A non-glacial interpretation for the Squantum Member is consistent with the results of previous studies by Dott (1961) and in most respects that of Socci and Smith (1987), though the latter favoured a glacially influenced marine setting on the basis of supposed ice-rafted dropstones in argillites. A glacial model is long standing (e.g., Sayles and LaForge, 1910;

Billings et al., 1939) and (as all theories do) reflects an earlier time in geological science when recognizing ancient glaciations and cold climates were based on limited criteria and understanding of sedimentary processes and deposits. The presence of poorly sorted

‗puddingstone‘ and ‗slate conglomerate‘ facies with finely laminated argillites were then seized by earlier workers as representing ancient tills and varves, respectively (Sayles, 1914, 1919;

Sayles and LaForge, 1910). Acceptance of a glacial origin of the Squantum Member was to have major ramifications in the broader field of geology. A generation of North American geologists rejected Wegener‘s arguments of continental drift (Wegener, 1929) until the late

1950‘s. Wegener argued that the famous American Squantum ‗Tillite‘ (then regarded as of

Permo-Carboniferous age) could not be glacial since it did fit into any glacial climate belt when placed on a reconstruction of Gondwana (Dott, 1961; Eyles, 2008, p. 160-161).

Wegener‘s reconstruction and entire thesis was rejected. Not until the 1960‘s were the Boston

Basin outcrops re-examined in the light of a better understanding of global sedimentary

47 processes and shown to be non-glacial (Dott, 1961) but also recognized as Neoproterozoic in age (Lenk et al., 1982).

In 1952, J. Crowell, who was very familiar with the pebbly mudstone deposits of

Californian mélanges produced by submarine slumping, concluded that ‗the case for a glacial origin (of the Squantum) was suspect if not downright wrong‘ (Pettijohn, 1984, p. 131; Eyles,

2008, p. 161). Pettijohn (1984, p. 128) applied Kuenen and Miglorini‘s classic 1950 paper on rhythmically laminated graded turbidites to further reinforce a non-glacial reinterpretation for the Squantum ‗varvites‘. Crowell (1957) was inspired to develop a facies model for the formation of a wide variety of diamictites by mass flow. Other workers developed non-genetic descriptors for diamictites that would enable objective descriptions of deposits and which too provided insights into the climatic significance of many ancient successions (e.g., ‗mixtite‘ and

‗debrite‘; Schermerhorn 1966, 1974; Eyles et al., 1983).

The study of Dott (1961) on the Squantum Member is demonstrably the earliest sophisticated application of newly developed facies investigations to the study of ancient glacial deposits; he showed that the Squantum Member bore the hallmark of a mass flow of rapidly deposited sediment in deep water and argued that there was a need for a similar re- evaluation of other ancient deposits. While accepting his conclusions, it is possible to take issue with one point explicit in the title of his paper. ‗Mass flow‘ and ‗glacial‘ are not exclusive processes and that it is not a case of either/or. Very commonly in glaciomarine settings, primary glacial debris generated along basin margins is reworked downslope. Such processes occur readily in ancient basins where the sedimentary record of landward environments is not often preserved. Nonetheless, such deposits still exhibit ancillary evidence of a glacially influenced setting in the form of glacially shaped ‗bullet‘ or ‗flat-iron‘ clasts, striations (though

48 these are easily destroyed), isolated large ‗dropstones‘ (or clusters, as bergs roll over) that are glacially shaped and striated. Compounding evidence may take the form of sedimentary evidence for repeated glacio-eustatic or isostatic changes across the basin, ice scour marks left by ice bergs and larger ‗glaciotectonic‘ deformation structures resulting from ice sheets and shelves grounding on the sea floor in deep water. Ice-contact glaciomarine successions in particular are extremely complex (see Benn and Evans, 2010) and highly variable stratigraphically and sedimentologically recording short-term changes in ice frontal position, sudden high-energy influxes of meltwater, rapid changes in water depths and sediment supply.

In contrast, the diamictites of the Boston Basin are locally generated lenses of diamictite intimately interbedded with ‗end member‘ facies consisting of clast-supported conglomerates and muddy turbidites.

Despite the work of Crowell (1957), Dott (1961), and the major review by

Schermerhorn (1974), the Squantum Member has recently been reported by Passchier and

Erukanure (2010) as the product of regional wet-based (or temperate) glaciation (glaciomarine diamictites deposited on a shelf or slope from floating ice-bergs). Their interpretation is based heavily on the association of diamictites with laminated fine-grained facies containing

‗outsized clasts‘, which they interpret to be ice-rafted dropstones (an alternate non-glacial explanation for these outsized clasts in presented in Chapter 3). Remarkably, their descriptions of the Squantum Member are limited to their examination of the diamictites exposed only at the type section. Significantly, Passchier and Erukanure (2010) explicitly acknowledge the lack of striated clasts and make note of the sedimentological evidence suggestive of a deep marine setting where mass wasting/debris flow activity occurred on a fan/slope surface. Another key result of Passchier and Erukanure (2010) study, which provides credence to the non-glacial

49 interpretation of the Squantum Member, are the chemical index of alteration (CIA) values they obtained from the matrix of the diamictites and finer grained facies at the type section which proved to be inconsistent with the scenario of large-scale continental glaciation or sea ice cover. The CIA values for the Squantum are considerably higher than typical Pleistocene glacial diamictites and they indicate that the chemical weathering regime for the diamictites requires significant exposure (ice-free conditions) of land during deposition.

Other diamictites linked to the Gaskiers glaciation are those of the Konnarock

Formation in southwestern Virginia (Miller, 1994) (formally the upper Mount Rogers

Formation; Jonas and Stose, 1939; King and Ferguson, 1960), the Grandfather Mountain

Formation of the southern Appalachians in eastern USA (Rankin, 1975) and the Granville

Formation of northern France (Dupret, 1974). These successions similarly consist of thick deep marine diamictite-turbidite successions preserved in arc-related basins within the Avalonian-

Cadomian Terranes (Rast et al., 1976; Williams & Hatcher, 1982, 1983; O‘Brien et al., 1983;

Williams, 1984). It has been suggested that the diamictites and conglomerates of the

Konnarock Formation records a glacially influenced ice-contact lake during repeated episodes of alpine glacial advance and retreat (Miller, 1994) but there is no unequivocal evidence of glaciation (Schwab, 1981). The Grandfather Mountain Formation of the southern Appalachians consists of diamictite and laminated pebble-cobble-boulder mudstone horizons that have been interpreted as tillites and varvites deposited in a deep lake or marine basin (Rankin, 1967;

Schwab, 1976). Alternatively, it has also been interpreted as alluvial-fan/braided river deposits

(Schwab, 1981) recording the reworking of fluvial sediments into a marine basin by mass flow

(Eyles, 1993). The Granville Formation consists of several kilometers of turbidites and volcanogenic rocks (Winterer, 1963). Just like the Squantum Member, Granville diamictites

50 record intermixing of conglomerates and mudstones (Winterer, 1963; Eyles, 1989). Winterer

(1963) showed that faintly striated clasts in the Granville Formation were produced by differential motion of clasts within a sandy matrix.

In the Snowball Earth model, the Squantum Member is correlated with the Gaskiers

Formation in Newfoundland (Williams and King, 1979; Eyles and Eyles, 1989) and is used to validate the occurrence of a global or near-global glaciation at c. 580 Ma (Gaskiers glaciation).

However, the lack of evidence for widespread glaciation along the Avalonian-Cadomian

Terranes indicates that the Gaskiers glaciation was limited in extent, recording the growth of local alpine glaciers on high volcanic peaks along the Avalon Terrane that delivered sediment and water into a deep marine basin, perhaps in the form of lahars. A modern analogue might well be the small glaciers on the summits of large active stratovolcanoes strung along chain of volcanic arcs (e.g., the Aleutian chain in Alaska; see Molnia, 2008).

An appropriate depositional analogue for these successions is modern volcanic arc systems with related arc basins such as the present-day Lesser Antilles Arc and adjacent

Grenada Basin in the Caribbean (see Chapter 5). In these settings large volumes of volcanic and sedimentary debris are supplied to marine basins flanking eruptive centers by debris flows, slumping and sliding due to seismic activity, high deposition rates, steep slopes and water saturated sediment (Carey and Sigurdsson, 1984; Fisher and Schmincke, 1994). The resulting basinal products are hemipelagic, clay-rich volcanic silt interbedded with volcaniclastic turbidites, conglomerates, and diamictites (Sigurdsson et al. 1980; Fisher, 1984; Deplus et al.,

2001; Picard et al., 2006). In summary, there is no compelling evidence for a widespread

Gaskiers glaciation.

Ironically, it is very interesting to note that other Neoproterozoic (c. 595 to 630 Ma) submarine-slide and ‗tillite-like‘ debris flow deposits present in other arc basins of the Avalon

Terrane were never interpreted as ‗glacial‘ (Skehan et al., 1985; Bailey et al., 1989). These deposits are characterized by large olistoliths of quartzarenite and carbonate, thick chaotic diamictite deposits (mélanges; Bailey et al., 1989), mud-rich and fine sand turbidites, extensive soft-sediment and syn-depositional deformation features indicative of subaqueous gravity mass transport. Slump-folded intervals are common and each sequence is associated with and interbedded by mafic and felsic volcanic rocks. These olistostrome deposits include the

Westboro Formation (1.1 km thick) located north and west of Boston, the Blackstone Group

(~2 km thick) of northern Rhode Island, and the Newport Neck Formation (0.5 to 1.3 km thick) of southern Narragansett Bay (Fig. 2.13). Deposition occurred in a slope/fan complex in a back-arc basin during the initial rifting stages of the Avalon Terrane and Gondwana at c. 630

Ma (Newport-Westboro-Blackstone Back-arc Basin; Bailey et al., 1989). These were derived from the failure of the steep fault-bounded margins of a plutonic/ volcanic source terrane (West

African or Amazonian cratonal sequence and Avalon Terrane) during crustal extension and basin rifting (Bailey et al., 1989).

A detailed description of these mass flow facies is beyond the scope of this study but it is important to make note of them as they provide further evidence for widespread episodes of rifting and mass flow contemporaneous with sedimentation in the Boston Basin (see Chapter 5 of this thesis for further discussion). The exposure of these olistostrome facies in laterally extensive outcrops provided the necessary contextual evidence for a non-glacial mass flow origin. In contrast, the more limited exposure of strata within the Boston Basin favoured a

52 much more restricted approach based simply on the outcrop similarity of laminated pebbly argillites with glaciolacustrine varves and diamictites with modern tills (Sayles, 1914).

The results presented above have major implications for paleoclimate models and understanding of the evolution of the biosphere. A non-glacial origin for the Squantum

Member combined with the broader application of a mass flow model to the fills of other arc- related circum-North Atlantic basins, undermines the notion of a globally widespread

Gaskiers-aged Snowball Earth event. This agrees with an emerging consensus that the Gaskiers glaciation was strictly regional in scope (Eyles and Eyles, 1983; Condon et al., 2002; Eyles and

Januszczak, 2004; Stern et al., 2006; Eyles, 2008; Zhao et al., 2009; Zhao and Zheng, 2010).

The widespread presence of diamictites within the Avalonian-Cadomian Terranes at this time is a function not of climate but tectonics and the availability of large volumes of sediment to rapidly subsiding fault-bounded arc basins. Narbonne and Gehling (2003) argued that the short-lived (c. 580 Ma) Gaskiers glaciation was a Snowball Earth event that had a major effect on the marine biosphere ultimately triggering the Cambrian ‗explosion.‘ This model is considered flawed in several important respects. First, a Snowball Earth model is not supported by facies data and reconstructions of depositional environments (see above). Second, given the restricted regional extent of ice covers during the Gaskiers glaciation it is considered very improbable that climate cooling had any marked effect on oceanic environments. Third, new work by paleobiologists shows that the Cambrian diversification of skeletal animals began much later at 540 Ma (close to the boundary of the Lower Cambrian at 542 Ma) and was much more gradual than previously thought occurring over some 20 million years (see Maloof et al.,

2010). Rifting of the supercontinent Rodinia, and attendant fundamental paleogeographic re- organizations with marked changes in geochemical cycles and sea level are now seen as more

53 viable ‗uniformitarian‘ explanations. It is likely too, that glaciation is also a response to supercontinent rifting at this time and the generation of high topography rather than catastrophic global climate cooling (Eyles, 2008).

2.8 CONCLUSIONS

Sedimentological investigations reported here from the Squantum Member of the

Boston Basin, USA, indicate that there is no case to be made for the continued usage of the term Squantum ‗Tillite‘. As argued originally by Dott (1961), diamictites are non-glacial debrites (‗mixtites‘; Schermerhorn, 1974) produced by the subaqueous slumping and mixing of conglomerates and fine sediment. These facies belong to a wider family of mass flow facies present within the Boston Basin such as clast- and matrix-supported conglomerates and sandstones deposited within the context of a rapidly subsiding, volcanically influenced and turbidite-dominated deep water marine basin. A non-glacial origin for the Squantum Member combined with the broader application of this depositional model to the fills of other arc- related circum-North Atlantic basins, undermines any notion of a globally widespread Gaskiers

Snowball Earth event. The widespread occurrence of diamictites (debrites) within the

Avalonian-Cadomian Terranes at this time is a function not of a glacial climate but tectonics and the availability of large volumes of sediment to rapidly subsiding, fault-bounded deep water arc-related basins.

Figure 2.1 Reconstruction of Gondwana at c. 570 Ma showing paleogeographic position of the Neoproterozoic island arc Avalonian-Cadomian Terranes (peri-Gondwana terranes) (modified from Balintoni et al., 2010).

Figure 2.2 (A) Location of the Boston Basin in eastern Massachusetts, USA, showing location of type locality of the Squantum Member at Squantum Head (Quincy, MA) (modified from Socci & Smith, 1987).

Figure 2.2 (B) Simplified geology of the Boston Basin (outlined in dashed line) showing location of study sites. 1: Squantum Head; 2: Hewitts Cove; 3: World‘s End Reservation Area; 4: Nantasket Avenue; 5: Nantasket Beach; 6: Franklin Park; 7: New Calvary Cemetery; 8: Arnold Arboretum; 9: Nahanton Park; 10: Hammond Pond Parkway Conservation Area; 11: Hyde Park; 12: Webster Conservation Area; 13: Chestnut Hill Mall; 14: Hemlock Gorge; 15: Mystic River Quarry. For regional geological setting see Figure 2.11.

Figure 2.3 (A) Clast-supported pebble-to-boulder conglomerate with clasts exhibiting a strong preferred orientation (site 3, Fig. 2.2B/ see Fig. 2.8 E for corresponding log). (B) Clast-supported cobble-to-granule conglomerate partially mixed with laminated argillite facies showing soft-sediment deformation (site 1, Fig. 2.2B).

Figure 2.3 (C) Enlarged section of partially mixed cobble-to-granule conglomerate and laminated argillite facies shown in Figure 2.3B (site 1, Fig. 2.2B). (D) Matrix- supported pebble-cobble conglomerate supported by coarse sand matrix (site 13, Fig. 2.2B).

Figure 2.3 (E) Matrix-supported pebble-cobble conglomerate with massive coarse- grained sandstone lenses (site 13, Fig. 2.2B/ see Fig. 2.8 F for corresponding log). (F) Crudely stratified succession of pebble conglomerate with beds of coarse-grained granule-sandstone (site 8, Fig. 2.2B/ see Fig. 2.8 G for corresponding log).

Figure 2.4 (A) Massive, fine-grained sandstone with deformed coarse-grained sandstone lenses or ‗wisps‘ (site 1, Fig. 2.2B). (B) Low-angle cross-laminated medium-to fine-grained sandstone (site 13; Fig. 2.2B).

Figure 2.4 (C) Laminated and normally graded sandstone (site 1, Fig. 2.2B/ see Fig. 2.8I for corresponding log). (D) Succession of laminated and normally graded sandstone deformed by large overturned fold (site 1, Fig. 2.2B).

Figure 2.4 (E) Deformed laminated and normally graded sandstone showing pillow structures (site12, Fig. 2.2B). (F) Medium-to-fine-grained massive sandstone displaying ‗dish-and-pillar‘ water-escape features (outlined in dashed line) (site 4, Fig. 2.2B).

Figure 2.5 (A) Matrix-supported, homogeneous diamictite (site 1; Fig. 2.2B). (B) Clast-supported diamictite showing crude grading, passing upwards into massive matrix-supported diamictite (site 1; Fig. 2.2B).

Figure 2.5 (C) Heterogeneous diamictite showing chaotically mixed clast-supported conglomerate with muddy-siltstone (site 1, Fig. 2.2B/ see Fig. 2.8 D for corresponding log). (D) Heterogeneous diamictite showing incomplete mixing of matrix-rich diamictite withclast-rich diamictites (site 1, Fig. 2.2B/ see Fig. 2.8 C for corresponding log).

Figure 2.5 (E) Rounded, massive mudstone slab within matrix-supported, homogeneous diamictite (site1; Fig. 2.2B). (F) Laminated and normally graded argillite (site 1; Fig. 2.2B).

Log C Log A Log B

B Laminated pebbly argillite

Laminated argillite

Figure 2.6 (A) Diamictite at type section at Squantum Head interbedded with a thin succession (cumulative thickness of 52 cm) of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies. The locations of stratigraphic logs shown in Figures 2.8A-C are labeled (logs A-C) (site 1; Fig. 2.2B). (B) Enlarged section of an alternating sequence of laminated, normally graded argillite and laminated pebbly argillite facies at Squantum Head; contact between two facies indicated by dashed line (site 1; Fig. 2.2B).

Figure 2.6 (C) Cut-slab of alternating sequences of laminated, normally graded argillite and laminated pebbly argillite facies exposed at Squantum Head (site 1; Fig. 2.2B). Length of slab is 20 cm. Arrow indicates ‗way-up‘ direction.

Figure 2.7 (A) Beds of reworked lapilli tuff (dark grey coarse-grained beds) within a succession of laminated, normally graded argillite (purple-grey laminated rocks) exposed along the southeast coast of Squantum Head (site 1; Fig. 2.2B) (see Figure 2.8H for corresponding log). (B) Massive bed of reworked lapilli tuff (site 1; Fig. 2.2B) (refer to Figure 2.7 A).

Scale |:------| 400 Microns

Figure 2.7 (C) Photomicrograph of lapilli tuff showing subangular to subrounded clasts quartz, iron oxides, alkali feldspar, and fresh sodic plagioclase, abundant fine felsic (rhyolite) tuffs, and rounded volcanic clasts of andesite and dacite. (D) Photomicrograph of laminated, normally graded argillite consisting of subangular-subrounded silt-sized particles of quartz, plagioclase, feldspar, melaphyre and dark grains of volcanic glass.

C are from the type locality (site 1; Fig. 2.2B; also see see also 2.2B; Fig. 1; (site locality type the from are C

- mined in study srea. Logs A Logs srea. study in mined

Stratigraphic logs of key outcrops exa outcrops key of logs Stratigraphic location). for 2.6A Figure

Figure 2.8 Figure

sandstone

interbed supported

conglomerate

Massive, matrix

Massive

C M S F G Cg

1 m 1

0.5 m 0.5

2.2B). (Fig. 13 site F: and 6, site E: 1, site D: .

supported

conglomerate

Massive, clast

C S F C CgM G

1 m 1 3 m 3 m 2

Stratigraphic logs of key outcrops examined in study area study in examined outcrops key of logs Stratigraphic

2.8 Figure

. G: site 8; H: site 1; I: site 1 (Fig. 2.2B). (Fig. 1 site I: 1; site H: 8; site G: .

rops examined in study area study in examined rops

Stratigraphic logs of key outc key of logs Stratigraphic

Figure 2.8 Figure

Figure 2.9 A 2.5 m-thick clast-supported conglomerate bed resting on (at base of hammer) a pyroclastic flow deposit (Brighton Volcanics) (site 3; Fig. 2.2B).

C D E

Figure 2.11 (A) A thick succession (16.5 m thick) of massive, clast-supported conglomerate conformably overlying heterogeneous diamictites, which in turn conformably overlies heavily-deformed massive argillite (site 1; Fig. 2.2B). (B) Corresponding stratigraphic log of outcrop shown in Figure 2.10A, highlighting the lithologic gradations that exist between conglomerate, diamictites and argillite facies at this section. (C) Heterogeneous diamictites underlying conglomerate units. (D) Contact (dashed line) between heterogeneous diamictites and lower massive and deformed argillite. (E) Lower-most unit of massive and deformed argillite with sandstone stringers.

Figure 2.12 Schematic of inferred paleoenvironment of the Boston Basin between c. 595-570 Ma with cross-section through the basin fill showing inferred stratigraphic relationships and depositional mechanisms of the lithofacies of the Boston Bay Group.

Area Enlarged

Atlantic Ocean

Figure 2.13 Map of eastern USA showing location of the ‗Avalonian Olistostromes‘ in relation to the Boston Basin. (A) Westboro Formation located north and west of the Boston Basin, (B) Blackstone Group in northern Rhode Island, (C) the Newport Neck Formation in southern Rhode Island (modified from Bailey et al., 1989). 78

REFERENCES

Akdağ, K. & Kırmaci, M. Z. (2008). Submarine fan and slope-apron deposition in a Cretaceous Forearc Basin: The Gürsökü Formation (Kavak–Samsun, N. Turkey). Journal of Asian Earth Sciences, 31(4-6), 429-438.

Allen, P. A. & Etienne, J.L. (2008). Sedimentary challenge to Snowball Earth. Nature Geoscience, 1, 817.

Andrews, J. T., Kirby, M., Jennings, A. E., & Barber, D. C. (1998). Late Quaternary stratigraphy, chronology, and depositional processes on the slope of S. E. Baffin Island, detrital carbonate and Heinrich Events: Implications for onshore glacial history. G´eographie Physique et Quaternaire 52, 1– 15.

Ault, A. (2003). Geology of Worlds End, Hingham, Massachusetts: sedimentation and volcanism within the Avalonian magmatic arc. Senior thesis, Wellesley College, Massachusetts, 100 pp.

Ault, A., Thompson, M., Grunow, A., & Ramezani, J. (2004). Mafic Mattapan volcanism at World‘s End, Hingham, Massachusetts: New dimensions of arc activity in the Southeastern New England Avalon Zone. Geological Society of America Abstracts with Programs, 36(2), 130.

Ayres, L.D., Van Wagoner, N.A., & Ferreira, W.S. (1991). Voluminous shallow-water to emergent phreatomagnetic basaltic volcaniclastic rocks, Proterozoic (~1886 MA) Amisk Lake Composite Volcano, in the Flon Greenstone Belt, Canada. Sedimentation in Volcanic Settings: SEPM Special Publication, 45, 175-187.

Bailey, R.H., Newman, W. A. & Genes, A. (1976). Geology of the Squantum "Tillite". In: B. Cameron (Ed.), Geology of Southern New England, 68th Annual Meeting New England Intercollegiate Geological Conference, Science Press, Princeton, NJ, 92-106.

Bailey, R. H. (1984). Origin of dropstones in the Precambrian Boston Bay Group. Geological Society of America Abstracts with Programs, 16(2), 2.

Bailey, R. H. (1987). Stratigraphy of the Boston Bay Group, Boston Area, MA. In: D.C. Roy (Ed.), Northeastern Section of the Geological Society of America: Boulder, Colorado, Geological Society of America, Centennial Field Guide, 5, 209-212.

Bailey, R.H., Skehan, S.W., Dreier, R.B., & Webster, M.J. (1989). Olistostromes of the Avalonian terrane of southeastern New England. Intercollegiate Geological Conference Meeting Guidebook, 145-164.

Bailey, R.H. & Bland, B.H. (2000). Ediacaran fossils from the Neoproterozoic Boston Bay Group, Boston area, Massachusetts. Geological Society of America Abstracts with Programs, 32, A11.

Bailey, R. H. (2005). Stratigraphy and depositional setting of Neoproterozoic strata on Southern Boston Harbor Islands, Boston Bay Group, Massachusetts. Geological Society of America, Northeast Section (Abstract), 27(7), 136.

Balintoni, C., Ducea, M.N, Zaharia, L., Chen, F, M. Clive, M., Hann, H.P., Li-Qui, L., & Ghergari, L. (2010). Late Cambrian–Ordovician northeastern Gondwanan terranes in the basement of the Apuseni Mountains, Romania. Journal of the Geological Society, London, 167, 1131–1145.

Barker, S.P., Haughton, P.D.W., McCaffrey, W.D., Archer, S.G. & Hakes, B. (2008). Development of rheological heterogeneity in clay-rich high-density turbidity currents: Aptian Britannia Sandstone Member, U.K. continental shelf. Journal of Sedimentary Research, 78, 45–68.

Bass, J.H. (2004). Conditions for formation of massive turbiditic sandstones by primary depositional processes. Sedimentary Geology, 166, 293–310.

Bell, K.G. (1948). Geology of the Boston metropolitan area. Ph.D thesis, Cambridge, Massachusetts Institute of Technology, 397 p.

Benn, D.I. & Evans D.J.A (2010). Glaciers and Glaciation. 2nd Edition, London, Hodder Education.

Benvenuti, M. (2003). Facies analysis and tectonic significance of lacustrine fan deltaic successions in the Pliocene–Pleistocene Mugello Basin, Central Italy. Sedimentary Geology, 157, 197– 234.

Billings, M.P. (1929). Structural geology of the eastern part of the Boston Basin. American Journal of Science, 5(18), 99-137.

Billings, M.P., Loomis, F.B., Jr., & Stewart, G.W. (1939). Carboniferous topography in the vicinity of Boston, Massachusetts. Bulletin: Geological Society of America, 50, 1867-1884.

Billings, M.P., & Tierney, F.L. (1964). Geology of the City Tunnel Extension. Journal of the Boston Society of Civil Engineers, 51, 111-154.

Billings, M.P. (1976). Geology of the Boston Basin. In: J.W. Skehan, and D.P. Murray (Eds.), Studies in New England Geology. Memoir: Geological Society of America, 46, 5-30.

Billings, M.P. (1979). Bedrock geology of the Boston Basin. In: Carboniferous Basins in Southeastern New England, Field Guidebook for trip No. 5 (Eds.), Cameron, B. 9th International Congress of Carboniferous Stratigraphy and Geology, 46-64.

Bingen, B., Griffin, W.L., Torsvik, T.H. & Saeed, A. (2005). Timing of Late Neoproterozoic glaciation on Baltica constrained by detrital zircon in the Hedmark Group, southeast Norway. Terra Nova, 17, 250-258.

Bond, G.C. & Lotti, R. (1995). Iceberg discharges into the North Atlantic on millennial time scales during the last glaciation. Cience, 267, 1005-1010.

Boswell, P.G. H. (1948). The thixotropy of certain sedimentary rocks. Science Progress, 36, 412-422.

Bouma, A. H. (1962). Sedimentology of some flysch deposits, a graphic approach to facies interpretation. Amsterdam: Elsevier

Bowman, D. Andrey, K., & Porat, N. (2004). Late-Pleistocene seismites from Lake Issyk-Kul, Tien Shan range, Kyrghyzstan. Sedimentary Geology, 163, 211 –228.

Brandes, C., Astorga, A., Littke, R., & Winsemann, J. (2008). Basin modelling of the Limon Back-arc Basin (Costa Rica): burial history and temperature evolution of an island arc-related basin-system. Basin Research, 20,119–142.

Busby, C.J. & Ingersoll, R.V.(1995). Tectonics of Sedimentary Basins. Cambridge, Massachusettes, USA: Blackwell Science.

Butterfield, N. J. (2007). Macroevolution and macroecology through deep time. Palaeontology, 50(1), 41–55.

Caldwell, D.W. (1964). The Squantum Formation: Paleozoic tillite or tilloid? In: J.W. Skehan (Ed.), Guidebook to Field Trips in the Boston Area and Vicinity. New England Intercollegiate Geological Conference, 56th Annual Meeting, 53-60.

Cameron, B. (1979). Evidence for glaciation in the Boston Bay Group. In: B. Cameron (Ed.), Carboniferous Basins of Southeastern New England. Field guidebook fro trip No. 5, 9th International Congress of Carboniferous Stratigraphy and Geology, 75-80.

Cardoza, K.D., Hepburn, J.C., & Hon, R. (1990). Geochemical constraints on the paleotectonic settings of two late Proterozoic mafic volcanic suites, Boston-Avalon Zone, Eastern Massachusetts. In: A.D. Socci, J.W Skehan, & G.W. Smith (Eds.), Geology of the composite Avalon terrane of southern New England. Geological Society of America Special Paper, 245, 113-131.

Carey, S., & Sigurdsson, H. (1984). A model of volcanogenic sedimentation in marginal basins. In: B. P. Kokelaar, & M. F. Howells (Eds.), Margin Basin Geology. Geological Society of London, 37-58.

Carto, S. L. & Eyles, N. (2011). The Squantum Member of the Boston Basin, Massachusetts, USA. In: Arnaud, E., Halverson, G. P. & Shields-Zhou, G.A. (Eds.), The Geological Record of Neoproterozoic glaciations. Geological Society of London, Memoir, (in press).

Chough, S.K., Hwang, I.G., & Choe, M.Y. (1990). The Miocene Doumsan Fan-Delta, southeast Korea: a composite fan-delta system in back-arc margin. Journal of Sedimentary Petrology, 60, 445-455.

Chun, S.S., Choe, M.Y., & Chough, S.K. (2002). Armored mudstone boulders in submarine debris-flow deposits, the Hunghae Formation, Pohang Basin: An evidence for the large-scale slumping of adjacent area of a submarine channel or scar wall. Geosciences Journal, 6(3), 215- 225.

Condon, D.J., Prave, A.R., & Benn, D.I. (2002). Neoproterozoic glacial-rainout intervals: Observations and implications. Geology, 30(1), 35–38.

Corsetti, F.A, Olcott, A.N., & Bakermans, C. (2006). The biotic response to Neoproterozoic snowball Earth. Palaeogeography, Palaeoclimatology, Palaeoecology, 232, 114-130.

Crosby, W.O. (1880). Contributions to the geology of eastern Massachusetts. Occasional Papers of the Boston Society of Natural History (III), 286 pp.

Crowell, J.C. (1957). Origin of pebbly mudstones. Bulletin: Geological Society of America, 68, 993-1010.

Crowell, J. C. (1964). Climatic significance of sedimentary deposits containing dispersed megaclasts, In: E.I.M. Nairn (Ed.), Problems in Paleoclimatology. Interscience Publishers, London, UK, 86-99 and 110-111.

Crowell, J.C. (1999). Pre-Mesozoic ice ages: their bearing on understanding the climate system. Memoir: Geological Society of America, 192, 106 pp.

Crowley, T.J., Hyde, W.T., & Peltier, W.R. (2001). CO2 levels required for deglaciation of a ‗near-snowball‘ Earth. Geophysical Research Letters, 28, 283–286.

De Blasioa, F.V., Elverhbia, A., Isslera, D., Harbitzb, C.B., Brync, P., Lienc R. (2004). Flow models of natural debris flows originating from overconsolidated clay materials. Marine Geology, 213, 439–455.

Deplus, C., LeFriant, A., Boudon, G., Komorowski, J.C., Villemant, B., Harfod, C., Segoufin, J., & Cheminee J.L. (2001). Submarine evidence for large-scale debris avalanches in the Lesser Antilles arc. Earth and Planetary Sciences Letters, 192, 145–157.

Deynoux. M., Ciner, A., Monod, O., Karabiyikoglu, M., Manatschal, G. & Tuzcu, S. (2005). Sedimentary Geology, 173, 315–343.

Donnadieu, Y., Fluteau, F., Ramstein, G., Ritz, C., & Besse, J. (2003). Is there a conflict between the Neoproterozoic glacial deposits and the snowball Earth interpretation: An improved understanding with numerical modelling. Earth and Planetary Science Letters, 208, 101–112.

Dott, R.H. Jr. (1961). Squantum ―tillite‖, Massachusetts—evidence of glaciation or subaqueous mass movements? Bulletin: Geological Society of America, 72, 1289-1305.

Dowdeswell, J.A., Whittington, R.J., Jennings, A.E., Andrews, J.T., Mackensen, A. & Marienfeld, P. (2000). An origin for laminated glacimarine sediments through sea-ice build-up and suppressed iceberg rafting. Sedimentology, 47, 557-576.

Dupret, L. (1974). Relations structurales entres les formations briovériennes de la region de Granville-saint Pair (Massif armoricain). Thése 3ocycle, Lab. Geol. Arm. Caen (inédit).

Durfee Cardoza, K., ms, (1987). Geochemistry of the Middlesex Fells Volcanics (mafic member) and Brighton Volcanics, Boston-Avalon zone, eastern Massachusetts. Master of Science Thesis, Boston College, Chestnut Hill, Massachusetts, 130 pp.

Einsele, G. (2000). Sedimentary Basins; Evolution, Facies, and Sediment Budget. Springer, Berlin, 792 pp.

Elverhbi, A., De Blasio, F.V., Butt, F.A., Issler, D., Harbitz, C.B., Engvik, L., Solheim, A., Marr, J. (2002). Submarine mass-wasting on glacially influenced continental slopes— processes and dynamics. In: Dowdeswell, J.A., O‘Cofaigh, C. (Eds.), Glacier-Influenced Sedimentation on High Latitude Continental Margins. Geological Society of London, London, UK, pp. 73– 87.

Emerson, B.K. (1917). Geology of Massachusetts and Rhode Island. US Geological Survey, Bulletin, 597, 289 pp.

Evans, D. A., Beukes, N. J., & Kirschvink, J. L. (1997). Low-latitude glaciation in the Palaeoproterozoic era. Nature, 386, 262–266.

Evans, A.D. (2000). Stratigraphic, geochronological, and palaeomagnetic constraints upon the Neoproterozoic climatic paradox. American Journal of Science, 300, 347–433.

Eyles, N., Eyles, C.H., & Miall, A.D. (1983). Lithofacies types and vertical profile models: an alternative approach to the description and environment interpretation of glacial diamict and diamictite sequences. Sedimentology, 30(3), 393-410.

Eyles, C.H., & Eyles, N. (1983). A glaciomarine model for Upper Precambrian diamictites of the Port Askaig Formation, Scotland. Geology, 11, 692-696.

Eyles, N., & Eyles, C.H. (1989). Glacially influenced deep-marine sedimentation of the Late Precambrian Gaskiers Formation, Newfoundland, Canada. Sedimentology, 36, 601-620.

Eyles, N. (1989). Tectonic controls on late Precambrian "tillite" sequences; examples from the Canadian Appalachians and the American Massif of France; Twenty-eighth international geological congress (abstracts). International Geological Congress, American Geological Institute Abstracts–Congress Geologique Internationale, 28(1), 467.

Eyles, N. (1990). Marine debris flows: Late Precambrian ‗‗tillites‘‘ of the Avalonian– Cadomian orogenic belt. Palaeogeography, Palaeoclimatology, Palaeoecology, 79, 73–98.

Eyles, N. (1993). Earth‘s glacial record and its tectonic setting. Earth-Science Reviews, 35, 1 – 248.

Eyles, C.H., Eyles, N., & Franca, A.B. (1993). Glaciation and tectonics in an active intracratonic basin: the Late Paleozoic Itarare Group, Parana Basin, Brazil. Sedimentology, 40, 1-25.

Eyles, C.H., & Eyles, N. (2000). Subaqueous mass flow origin for Lower Permian diamictites and associated facies of the Grant Group, Barbwire Terrace, Canning Basin, Western Australia. Sedimentology, 47, 343– 356.

Eyles, N., & Januszczak, N. (2004). ‗Zipper-rift‘: A tectonic model for Neoproterozoic glaciations during the break-up of Rodinia after 750 Ma. Earth Science Reviews, 65, 1-73.

Eyles, N. (2008) Glacioepochs and the supercontinent cycle after ~3.0 Ga: tectonic boundary conditions for global cooling. Palaeogeography, Palaeoecology, Palaeoclimatology, 258, 89- 129.

Fairchild, I. J., & Kennedy, M.J. (2007). Neoproterozoic glaciation in the Earth System Journal of the Geological Society, 164, 895–921.

Fischer, R.V., & Schminke, H.U. (1984). Pyroclastic Rocks. Springer-Verlag, Berlin.

Fisher, R.V. (1984). Submarine volcaniclastic rocks. Geological Society of London Special Publications, 16, 5-27.

Fisher, R.V., & Schminke, H.U. (1992). Volcaniclastic sedimentation and facies. In: K. Pye (Ed.) Seiment transport and depositional processes. Blackwell Scientific Publications, 351-388.

Flint, R.F., Sanders, J.E., Rodgers, J. (1960). Diamictite, a substitute term for symmictite. Geological Society of America Bulletin, 71, 1809–1810.

Folk, R.L. (1974). Petrology of Sedimentary Rocks. Hemphill, Austin, TX, 182 pp.

Frey-Martinez, J., Cartwright, J. & James, D. (2006). Frontally confined versus frontally emergent submarine landslides: a 3D seismic characterization. Marine Petrology and Geology, 23, 585–604.

Füchtbauer, H. (1988).Transportation procedures and sediment structures. In: Sediments and Sedimentary rocks, Part of 2: Sediment Petrology, (Ed.) H. Füchtbauer. Schweizerbart'sche Verlagsbuchhandlung (Nägele u. Obermiller), Stuttgart, 779-863.

Halverson, G., Hoffman, P.F., Schrag, D.P., Maloof, A.C., & Rice, A.H.N. (2005). Towards a Neoproterozoic composite carbon isotope record. Bulletin: Geological Society of America, 117, 1181–1207.

Hambrey, M.J., & Harland, W.B. (1981). Earth‘s Pre- Pleistocene Glacial Record. Cambridge: Cambridge University Press, 1004 pp.

Hampton, M. A. (1972). The role of subaqueous debris flow in generating turbidity currents. Journal of Sedimentary Petrology, 42, 745-793.

Hampton, M.A., Lee, H.J., & Locat, J. (1996). Submarine landslides. Review of Geophysics, 34, 33-59.

Harms, J.C., Southard, J.B., Spearing, D.R., & Walker, R.G. (1975). Depositional environments as interpreted from primary sedimentary structures and stratification sequences. Society of Economic Paleontologists and Mineralogists, Short Course 2, 161 pp.

Heinrich, H. (1988) Origin and consequences of cyclic ice rafting in the Northeast Atlantic Ocean during the past 130,000 years. Quaternary Research, 29, 142-152.

Hepburn, J.C., Hon, R., Dunning, G.R., Bailey, R.H., & Galli, K. (1993). The Avalon and Nashoba terranes (eastern margin of the Appalachian orogen in southeastern New England). In: J.T. Cheney, and J.C. Hepburn (Eds.), Field Trip Guidebook for the Northeastern United States. Geological Society of America, Annual Meeting, Boston, 2, X1-X31.

Hiscott, R.N. (1994).Traction-carpet stratification in turbidites- fact or fiction? Journal of Sedimentary Research, A64, 204-208.

Hoffman, P.F., Kaufman, A.J., Halverson, G.P., & Schrag, D.P. (1998). A Neoproterozoic Snowball Earth. Science, 281, 1342–1346.

Hoffman, P.F. & Schrag, D.P. (2000). Snowball Earth. Scientific American, 282, 62–75.

Hoffman, P. F. & Schrag D. P. (2002). The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14(3), 129-155.

Hoffman, P.F. (2005). On Cryogenian (Neoproterozoic) ice-sheet dynamics and the limitations of the glacial sedimentary record. South African Journal of Geology, 108, 557–577.

Hoffman, P.F. & Li, Z. X. (2009). A palaeogeographic context for Neoproterozoic glaciation Palaeogeography, Palaeoclimatology, Palaeoecology, 277, 158–172.

Hon, R. & Hepburn, J.C. (1986). Igneous geochemistry and its implications for terrane analysis of the Avalonian event in southeastern New England, U.S.A. Atlantic Geoscience Society Abstract, 1986 colloquium. Maritime sediments. Geology, 22, 331.

Hyde, W.T., Crowley, T.J., Baum, S.K., & Peltier, W.R. (2000). Neoproterozoic 'snowball Earth' simulations with a coupled climate/ice-sheet model. Nature, 405, 425–429.

Ilstad, T., De Blasio, F. V., Elverhøi, A., Harbitz, C.B., Engvik, L., Longva O. & Marr, J. (2004a). On the frontal dynamics and morphology of submarine debris flows. Marine Geology 213, 481-497.

Ilstad, T., Marr, J., Elverhøi, A., Issler, D. & Harbitz, C. (2004b). Laboratory studies of subaqueous debris flow by measurements of pore fluid pressure and total stress. Marine Geology, 213, 403-414.

Ilstad, T., Elverhøi, A., Issler, D. & Marr, J. (2004c). Subaqueous debris flow behavior and its dependence on the sand/clay ratio: a laboratory study using particle tracking. Marine Geology, 213, 415-438.

Ineson, J.R. (1989). Coarse-grained submarine fan and slope apron deposits in a Cretaceous back-arc basin, Antarctica. Sedimentology, 36, 793– 819.

Jonas, A.E., & Stose, G.W. (1939). Age relations of the Precambrian rocks in the Catoctin Mountain-Blue Ridge and Mount Rogers anticlinoria in Virginia. American Journal of Science, 237, 575-593.

Kaye, C.A., & Zartman, R.E. (1980). A Late Proteozoic Z to Cambrian age for the stratified rocks of the Boston Basin, Massachusetts. In: D.R. Wones (Ed.), The Caledonides in the USA: Blacksburg, Virginia Polytechnic Institute and State University Memoir, 2, 257-261.

Khoo, M. A., Ramezani, J., & Thompson, M. D. (2008). Age and Stratigraphic Affinity of Neoproterozoic Conglomerate at Neponset River Gorge, Milton, MA: Constraints from U-Pb Geochronology. Geological Society of America Abstracts with Programs, 40(6) 146.

King, P.B. & Ferguson, H.W. (1960). Geology of northeastern-most Tennessee. U.S. Geological Survey, Professional Papers, 311.Kirschvink, J.L. (2002). When All of the Oceans Were Frozen. Recherche, 355, 26–30.

Kirschvink, J.L. (2002). When All of the Oceans Were Frozen. Recherche, 355, 26–30.

Kneller, B.C. & Buckee, C. (2000). The structure and fluid mechanics of turbidity currents: a review of some recent studies and their geological implications. Sedimentology, 47, 62–94.

Knoll, A. H., Walter, M., Narbonne, G., & Christie-Blick N. (2006). The Ediacaran period: A new addition to the geologic time scale. Lethaia, 39, 13–30.

LaForge, L. (1932). Geology of the Boston area, Massachusetts. U.S. Geological Survey Bulletin, 839, 105.

Lahee, F.H. (1914). Late Paleozoic glaciation in the Boston Bain, Massachusetts. American Journal of Science, 187(4), 316-318.

Larsen, V. & Steel, R.J. (1978). The sedimentary history of a debris flow-dominated alluvial fan - a study of textural inversion. Sedimentology, 25, 37-59.

Leather, J., Allen, P., Brasier, M.D., & Cozzi, A. (2002). Neoproterozoic Snowball Earth under scrutiny: evidence from the Fiq glaciation of Oman. Geology, 30, 891– 894.

Lenk, C., Strother, P.K., Kaye, C.A., & Barghoorn, E.S. (1982). Precambrian age of the Boston Basin: new evidence from microfossils. Science, 216, 619-620.

Lindsay, J.F., Summerson, C., & Barrett, P. J. (1970). A long-axis clast fabric comparison of the Squantum 'tillite', Massachusetts and the Gowganda Formation, Ontario. Journal of Sedimentary Research, 40(1), 475-479.

Lowe, D.R., & LoPiccolo, R.D. (1974). The characteristics and origins of dish and pillar structures. Journal of Sedimentary Petrology, 44, 484– 501.

Lowe, D.R. (1982). Sediment gravity flows: II. Depositional models with special reference to the deposits of high-density turbidity currents. Journal of Sedimentary Petrology, 52, 279-297.

Lowe, D.R., & Guy, M. (2000). Slurry-flow deposits in the Britannia Formation (Lower Cretaceous), North Sea: a new perspective on the turbidity current and debris flow problem. Sedimentology, 47, 31–70.

MacGabhann, B. A. (2005). Snowball Earth, the Dalradian Supergroup, and the Cryogenian- Ediacaran Boundary. 48th Irish Geological Research Meeting Programme and Abstracts, 33.

Maloof, A.C., Porter, M. S., Moore, J.L., Dudas, F.O., Bowring, S.A., Higgins, J.A., Fike, D.A., Eddy, M.P. (2010). The earliest Cambrian record of animals and ocean geochemical change. Geological Survey of America Bulletin, 122(11/12), 1731-1774.

Martins-Neto, M.A. (1996). Lacustrine fan-deltaic sedimentation in a Proterozoic rift basin: the Sopa-Brumadinho Tectonosequence, southeastern Brazil. Sedimentary Geology, 106, 65-96.

Meert, J.G. (2001). Growing Gondwana and Rethinking Rodinia: A paleomagnetic perspective. Gondwana Research, 4(3), 279-288.

Meert, J.G. (2003). A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics, 362, 1-40.

Miall, A.D. (1978). Lithofacies types and vertical profile models in braided river deposits: a summary, In: A.D. Miall (Ed.), Fluvial Sedimentology. Memoir: Canadian Society of Petroleum Geologists, 5, 578- 586.

Miall, A.D. (1985). Sedimentation on an early Proterozoic continental margin under glacial influence: the Gowganda Formation (Huronian), Elliot Lake area, Ontario, Canada. Sedimentology, 32, 763–788.

Miller, J.M.G. (1994). The Neoproterozoic Konnarock Formation, southwestern Virginia, USA. Glaciolacustrine facies in a continental rift. In: M., Deynoux, J.M.G. Miller, E.W. Domack, N. Eyles, I.J. Fairchild & G.M. Young (Eds.) Earth‘s Glacial Record: World and Regional Geology 5, Cambridge University Press, 47-59.

Mohrig, D., Whipple, K.X., Hondzo, M., Ellis, C. & Parker, G. (1998). Hydroplanning of subaqueous debris flows. Geological Society of America Bulletin, 110, 387-394.

Mohrig, D., Elverhoi, A., & Parker, G. (1999). Experiments on the relative mobility of muddy subaqueous and subaerial debris flows, and their capacity to remobilize antecedent deposits. Marine Geology, 154, 117-129.

Mohrig, D. & Marr, J.G. (2003). Constraining the efficiency of turbidity current generation from submarine debris flows and slides using laboratory experiments. Marine Petrology and Geology, 20, 883–899.

Molnia, B.F. (2008). Glaciers of North America -- Glaciers of Alaska. In: R.S., Jr. Williams, & J.G. Ferrigno (Eds.), Satellite image atlas of glaciers of the world: U.S. Geological Survey Professional Paper 1386-K, 525 p.

Montenat, C., Barrier, P., Ott d′Estevou, P., Hibsch, C. (2007). Seismites: An attempt at critical analysis and classification. Sedimentary Geology, 196, 5–30.

Mulder, T., & Alexander, J. (2001). The physical character of subaqueous sedimentary density flows and their deposits. Sedimentology, 48, 269-299.

Mutti, E. (1992). Turbidite sandstones. Agip Special Publication, Milan.

Murphy, J.B., van Staal, C.R., & Keppie, J.D. (1999). Middle to Late Paleozoic Acadian Orogeny in the northern Appalachians: A Laramidestyle plume-modified orogeny. Geology, 27, 653-656.

Murphy, J.B., Fernandez-Suarez, J., Keppie, J.D., and Jeffries, T.E. (2004). Contiguous rather than discrete Paleozoic histories for the Avalon and Meguma terranes based on detrital zircon data. Geology, 32, 585-588.

Nance, R.D. (1990). Late Precambrian-early Paleozoic arc-platform transitions in Northern Appalachians; Review and implications. In: A.A. Socci, J.W. Skehan, & G.W. Smith (Eds.), Geology of the Composite Avalon terrane of Southern New England: Boulder, Colarado. Geological Society of America Special Paper, 245, 1-11.

Narbonne, G.M. (1998). The Ediacara biota: A terminal Proterozoic experiment in the evolution of life. Geological Survey of America Today, 8(2), 1-6.

Narbonne, G.M., & Gehling, J.G. (2003). Life after snowball: the oldest complex Ediacaran fossils. Geology, 31, 27-30.

Nemec, W., Steel, R.J., Porebski, S.J., & Spinnangr, A. (1984). Domba Conglomerate, Devonian, Norway: process and lateral variability in a mass flow-dominated lacustrine fan- delta. In: E.H. Koster, & R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Canadian Society of Petroleum Geologists, 295–330.

Nemec, W., & Steel, R.J. (1984). Alluvial and coastal conglomerates: their significant features and some comments on gravelly mass flow deposits. In: E.H. Koster, & R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Canadian Society of Petroleum Geologists, 1 – 31.

Nichols, G. (1999). Sedimentology & Stratigraphy. Blackwell Science, Osney Mead, Oxford. 62-75.

O‘Brien, S.J., Wardle, R.J., & King, A.F. (1983) The Avalon Zone: a Pan-African terrane in the Appalachian Orogen of Canada. Journal of Geology, 18, 195-222.

O‘Brien, S.J., O‘Brien, B.H., Dunning, G.R., & Tucker, R.D. (1996). Late Proterozoic and related peri-Gondwanan rocks of the Newfoundland Appalachians. In: R.D. Nance, & M.D. Thompson (Eds.), Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado. Geological Society of America Special Paper, 304, 9-28.

Oggiano, G. & Mameli, P. (2006). Diamictite and oolitic ironstones, a sedimentary association at Ordovician–Silurian transition in the north Gondwana margin: New evidence from the inner nappe of Sardinia Variscides (Italy). Gondwana Research, 9(4), 500-511.

Olcott, A.N., Sessions, A.L., Corsetti, F.A., Kaufman, A.J., & de Oliveria, T.F. (2005). Biomarker evidence for photosynthesis during Neoproterozoic Glaciation, Science, 310, 471- 474.

Passchier, S., & Erukanure, E. (2010). Paleoenvironments and weathering regime of the Neoproterozoic Squantum ―Tillite‖, Boston Basin: no evidence of a snowball Earth. Sedimentology, 57(6), 1526-1544.

Peltier, W.R. (2004). Global glacial isotasy and the surface of the ice-age Earth: The ICE-5G (VM2) Model and GRACE. Annual Reviews of Earth Planetary Science, 32, 111–149.

Pettijohn, F.J. (1957). Sedimentary rocks. New York, Harper and Brothers, 718 pp.

Pettijohn, F. J. (1984). Diary of an Unrepentant Field Geologist. University of Chicago Press, Chicago, 260 pp.

Picard, M., Schneider, J-C, & Boudon, G. (2006). Contrasting sedimentary processes along a convergent margin: the Lesser Antilles arc system. Geo-Marine Letters, 26 (6), 397-410.

Pickering, K., Dorrik, S., Watson, M., & HI Scott, R. (1986). Deep water facies, processes and models: A review and classification scheme for modern and ancient sediments. Earth-Science Reviews, 23, 75-174.

Pickering, K.T., Hi Scott, R.N., & Hein, F.J. (1989). Dee-marine environments: clastic sedimentation and tectonics. Unwin, London.

Pickering, K.T., & Corregidor, J. (2005). Mass-Transport Complexes (MTCs) and tectonic control on basin-floor submarine fans, Middle Eocene, south spanish Pyrenees. Journal of Sedimentary Research, 75, 761–783.

Pierrehumbert, R.T. (2004). High levels of atmospheric carbon dioxide necessary for the termination of global glaciation. Nature, 429 (6992), 646–649.

Piper, D.J.W., Cochonat, P., Ollier, G., Le Dreezen, E., Morrison, M., & Baltzer, A. (1992). Evolution progressive d'un glissement rotationnel en un courant de turbidite : cas du seÂisme de 1929 des Grand Bancs (Terre Neuve). CR Acad. Science Paris, 314 (Series II), 1057-1064.

Piper, D.J.W., Pe-Piper, G., & Ingram, S.C. (2004) Early Cretaceous sediment failure in the southwestern Sable Subbasin, offshore Nova Scotia. Bulletin: American Association of Petroleum Geologists, 88, 991–1006.

Prior, D.B., & Bornhold, B.D. (1990). The underwater development of Holocene fan deltas, In: A. Colella, & D.B. Prior (Ed.), Coarse-grained Deltas. International Association of Sedimentologists Special Publications, 10, 75-90.

Pyles, D.R., & Jeanette, D.C. (2009). Geometry and architectural associations of co-genetic debrite-turbidite beds in basin-margin strata, Carboniferous Ross Sandstone (Ireland Sedimentologists Special Publications, 10, 75-90.): Applications to reservoirs located on the margins of structurally confined submarine fans. Marine and Petroleum Geology, 26(10), 1974-1996.

Rankin, D.W., (1967). The stratigraphic and structural position of the upper Precambrian Mount Rogers volcanic group, Virginia, North Carolina and Tennessee (abstracts) Geological Society America, Southeastern Section, Program for Annual Meeting, 49p.

Rankin, G. (1975). The continental margin of eastern North America in the southern Appalachians : the opening and closing of the Proto-Atlantic Ocean. American Journal of Science, 275, 298-336.

Rast, N., O‘Brien, B. H. & Wardler, J. (1976). Relationships between Precambrian and Lower Palaezoic rocks of the ‗Avalon Platform‘ in New Brunswick, the northeast Appalachians and the British Isles. Tectonophysics, 30, 315-338.

Rast, N., & Skehan, J.W. (1983). The evolution of the Avalonian Plate. Tectonophysics, 100, 257-286.

Rice, H.N., Halverson, G.P., & Hoffman, P.F. (2003). Three for the Neoproterozoic: Sturtian, Marinoan and Varangerian glaciations. American Geological Union Joint Assembly, Nice, France.

Richards, M., Bowmanb, M., & Reading, H. (1998). Submarine fan systems I: characterization and stratigraphic prediction. Marine and Petroleum Geology, 15, 689-717.

Ross, W.C., Halliwell, B.A., May, J.A., Watts, D.E., & Syvitski, J.P.M. (1994). Slope readjustment: a new model for the development of submarine fans and aprons. Geology, 22, 511–514.

Rossetti, D.F. (1999). Soft-sediment deformation structures in late Albian to Cenomanian deposits, Sao Luis Basin, northern Brazil: evidence for palaeoseismicity. Sedimentology, 46, 1065–1081.

Samson, S. D., D'Lemos, R. S. Miller, B. V. & Hamilton, M. A. 2005. Neoproterozic palaeogeography of the Cadomian and Avalon terranes: constraints from detrital zircons U-Pb ages. Journal of the Geological Society, 1-7.

Sayles, R.W., & LaForge, L. (1910). The glacial origin of the Roxbury Conglomerate. Science, 32, 723-724.

Sayles, R.W. (1914). The Squantum Tillite. Bulletin: Harvard Museum Comparative Zoology, 66, 141-175.

Sayles, R.W. (1919). Seasonal deposition in the aqueo-glacial sediments. Memoirs of Museum Comparative Zoology, Harvard College, 47, 67 pp.

Schermerhorn, L.J.G. (1966). Terminology of mixed coarse-fine sediments. Journal of Sedimentary Petrology, 36, 831– 836.

Schermerhorn, L. J. G. (1974). Late Precambrian mixtites: Glacial and/or non-glacial? American Journal of Science, 274, 673–824.

Schermerhorn, J. L. G. (1976). The Aljustrel volcanics: megacryst tuff and green tuff (Aljustrel and Gavia˜o pyritite deposits, South Portugal). Memorias e Noticias Publicacoes¸ do Museo e Laboratorio Mineralogico da. Universidad de Coimbra, 82, 41–60.

Schermerhorn, L. J. G. (1983). Proterozoic glaciation in the light of CO2 depletion in the atmosphere. Memoirs: Geological Society of America, 161, 309– 315.

Schrag, D.P., Berner, R.A., Hoffman, P.F., & Halverson, G.P. (2002). On the initiation of a snowball Earth. Geochemical Geophysics Geosystems, 3, 1029-1036.

Schwab, F.L. (1976). Depositional environments, provenance, and tectonic framework; upper part of the late Precambrian Mount Rogers Formation, Blue Ridge Province, southwestern Virginia. Journal of Sedimentary Research, 46(1), 3-13.

Schwab, F.L. (1981). Late Precambrian tillites of the Appalachians, In: M.J. Hambrey, & W.B. Harland (Eds.), Earth‘s Pre-Pleistocene Glacial Record. Cambridge University Press, 751-755.

Sigurdsson, H., Sparks, R.S.J., Carey, S.N., & Huang, T.C. (1980). Volcanogenic sedimentation in the Lesser Antilles arc. Journal of Geology, 88, 523-540.

Skehan, S.J., & Murphy, D.P. (1980). Geological profile across southeastern New England. Tectonophysics, 69, 285-319.

Skehan, S.J., J.W., Bailey, R.H., Dreier, R.B., & Webster, M.J. (1985). Late Proterozoic Olistromes of the Avalon Zone, Southeastern Massachusetts and Rhode Island. Geological Society of America, Abstracts with Programs, 17(7) 719 pp.

Smith, G.W., & Socci, A.D. (1990). Late Precambrian sedimentary geology of the Boston Basin, In: A.A. Socci, S.J. Skehan, & G.W. Smith (Eds.), Geology of the composite Avalon terrane of southern New England. Geological Society of America Special paper, 245, 75-84.

Socci, A.D., & Smith, G.W. (1987). Recent sedimentological interpretations in the Avalon terrane of the Boston Basin, Massachusetts: Maritime Sediments and Atlantic. Geology, 23, 13-39.

Socci, A.D., & Smith, G.W. (1990). Stratigraphic implications of facies within the Boston Basin: In: A.D. Socci, J.W. Skehan, & G.W. Smith (Eds.), Geology of the Composite Avalon terrane of southern New England: Boulder, Colorado, Geological Society of America Special Paper, 245, 55-74.

Sohn, Y.K., Kim, S.B., Hwang, I.G., Bahk, J.J., Choe, M.Y., & Chough, S.K. (1997). Characteristics and depositional processes of large-scale gravelly Gilbert-type forsets in the Miocene Doumsan fan delta, Pohang basin, SE Korea. Journal of Sedimentary Research, 67, 130-141.

Sohn, Y.K. (2000a). Coarse-grained debris flow deposits in the Miocene fan deltas, SE Korea: a scaling analysis. Sedimentary Geology, 130, 45–64.

Sohn, Y.K. (2000b). Depositional processes of submarine debris flows in the Miocene fan deltas, Pohang Basin, SE Korea with special reference to flow transformation. Journal of Sedimentary Research, 70, 491–503.

Sohn, Y.K., Choe, M.Y., Jo, H. R. (2002). Transition from debris flow to hyperconcentrated flow in a submarine channel (the Cretaceous Cerro Toro Formation, southern Chile), Terra Nova, 14, 405–415.

Stow, D.A.V., Howell, D.G., & Nelson, C.H. (1985). Sedimentary, tectonic and sea-level controls. In: A. H. Bouma, N. E. Barnes & W. R. Normark (Eds.), Submarine Fans and Related Turbidite Sequences , 15-22. Springer- Verlag, New York.

Stow, D.A.V., & Johansson, M. (2000). Deep-water massive sands: nature, origin and hydrocarbon implications. Marine Petrology Geology, 17, 145-174.

Talling, P.J., Amy, L.A., Wynn, R.B., Peakall, J., & Robinson, M. (2004). Beds comprising debrite sandwiched within co-genetic turbidite: origin and widespread occurrence in distal depositional environments. Sedimentology, 51, 163– 194.

Taira, A. (2001). Tectonic Evolution of the Japanese Island Arc System. Annual Review Earth Planetary Science, 29,109–34.

Tierney, F.L., Billings, M.P., & Cassidy, M.M. (1968). Geology of the City Tunnel, Greater Boston, Massachusetts. Boston Society of Civil Engineers Journal, 55, 60-96.

Thompson, M.D. (1985). Evidence for a late Precambrian caldera in Boston, Massachusetts. Geology, 13, 641-643.

Thompson, M.D. (1993). Late Proterozoic stratigraphy and structure in the Avalonian magmatic arc southwest of Boston Massachusetts. American Journal of Science, 293, 725-753.

Thompson, M.D., Hermes, O.D., Bowring, S.A., Isachsen, C.E., Besancon, J.R., & Kelly, K.L. (1996). Tectonostratigraphic implications of Late Proterozoic U-Pb zircon ages in the Avalon Zone of southeastern New England. In: R.D. Nance and M.D. Thompson (Eds.), Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America Special Papers, 304, 179-191.

Thompson, M.D., & Bowring, S.A. (2000). Age of the Squantum ―tillite‖, Boston Basin, Massachusetts: U-Pb zircon constraints on terminal Neoproterozoic glaciation. American Journal of Science, 300, 630-655.

Thompson, M.D., Grunow, A.M., & Ramezani, J. (2003). New Paleomagnetic pole from the Late Cambrian Nahant Gabbro, Nahant, Massachusetts: implications for drift history of the Southeastern New England Avalon Zone. Geological Society of America Abstracts with Programs, 35, p. 18

Thompson, M.D., & Grunow, A.M. (2004). A Roxbury Review: conglomerates of greater Boston, Massachusetts. In: L.S. Hanson (Ed.), Guidebook to field trips in Rhode Island and Massachusetts: Salem, MA. 96th Annual Meeting of the New England Intercollegiate Field Conference, p. C2-1 - C2-21.

Thompson, M.D., Grunow, A.M., & Ramezani, J. (2007). Late Neoproterozoic paleogeography of the Southeastern New England Avalon Zone: Insights from U-Pb geochronology and paleomagnetism. Bulletin: Geological Society of America, 119, 681-696.

Torsvik, T.H. (2003). The Rodinia Jigsaw Puzzle. Science, 300, 1379-1381.

Tripsanas, E.K., Piper, D.J.W., Jenner, K. A., & Bryant, W. R. (2008) Submarine mass- transport facies: new perspectives on flow processes from cores on the eastern North American margin. Sedimentology, 55, 97–136.

Trop, J.M., Ridgway, K.D., Sweet, A.R., & Layer, P.W. (1999). Submarine fan deposystems and tectonics of a Late Cretaceous forearc basin along an accretionary convergent plate boundary, MacColl Ridge Formation, Wrangell Mountains, Alaska. Canadian Journal of Earth Science, 36, 433–458.

Walker, R.G. (1975). Generalized facies model for resedimented conglomerates of turbidite association. Bulletin: Geological Society of America, 86, 737–748.

Walker, R.G. (1984). Shelf and Shallow Marine Sands. In: R.G. Walker (Ed.), Facies Models (2nd edn), 141–170. Geological Association of Canada, Toronto.

Walker, R.G. (1992). Turbidites and submarine fans. In: R.G. Walker, & N.P. James (Ed.), Facies Models. Geological Association of Canada, Ontario, 239- 263.

Wegener, A. (1929). The origin of oceans and continents (4ed.). Braunschweig: Friiedwich, Vieweg and Sohn Akt. GES.

Weisner, M.G., Wang, Y., & Zheng, L. (1995). Fallout of volcanic ash to the deep South China Sea induced by the 1991 eruption of Mount Pinatubo (Philippines). Geology, 23, 885-888.

Williams, H., & King, A. F. (1979). Trepassey map area, Newfoundland. Canadian Geological Survey Memoirs, 389, 24 pp.

Williams, H. & Hatcher, R.D. JR. (1982) Suspect terranes and accretionary history of the Appalachian orogen. Geology, 10, 530-536.

Williams, H. & Hatcher, R.D. JR. (1983) Appalachian suspect terranes. In: R.D. Hatcher, Jr (Ed.) Contributions to the Tectonics and Geophysics of Mountain Chains. Geological Society of America, 158, 33-53.

Williams, H. (1984) Miogeoclines and suspect terranes of the Caledonian-Appalachian Orogen : tectonic patterns in the North Atlantic region. Canadian Journal of Earth Science, 21, 887- 901.

Winterer, E. (1963). Late Precambrian pebbly mudstone in Normandy, France. In: A.E.M. Nairn (Ed.), Problems in Palaeoclimatology, 159-178. Interscience, London

Vidal, G. (1976). Late Precambrian microfossils from the Visingso beds in southern Sweden. Fossils and Strata, 9, 1–57.

Zartman, R.E., & Naylor, R.S. (1984). Structural implications of some radiometric ages of igneous rock sin southeastern Massachusetts. Bulletin: Geological Society of America, 95, 522- 539.

Zen, E. (1983). Bedrock geologic map of Massachusetts: Reston, Virginia, U.S. Geological Survey, 3 sheeta, scale 1:250,000.

Zhao Y.Y., Zheng Y.F., & Chen F.K. (2009) Trace element and strontium isotope constraints on sedimentary environment of Ediacaran carbonates in southern Anhui, South China. Chemical Geology, 265, 345-362.

Zhao Y. Y., & Zheng Y.F. (2010) Stable isotope evidence for involvement of de-glacial meltwater in Ediacaran carbonates in South China. Chemical Geology, 271, 86-100.

CHAPTER 3:

REINTERPRETATION OF ‘ICE-RAFTED’ PEBBLY ARGILLITES OF THE NEOPROTEROZOIC SQUANTUM MEMBER (BOSTON BASIN, USA) AS NON- GLACIAL DEBRITE-TURBIDITE COUPLETS

ABSTRACT

The exceptionally thick (~7 km) Neoproterozoic volcanic-sedimentary fill of the

Boston Basin (c. 593-570 Ma) in eastern Massachusetts, USA, comprises interbedded marine conglomerate, diamictite, sandstone, and argillite facies. These lithofacies record the delivery of large volumes of volcaniclastic sediment by widespread mass flow to deep water fans within a complex faulted and volcanically influenced arc basin at the mid-latitude margin of

Gondwana that developed between c. 630-570 Ma. Some 100 years ago, the diamictites

(Squantum Member) and the 5.5-km-thick succession of laminated argillites (Cambridge

Formation) of the Boston Basin were attributed to glacial processes (e.g., glacial tillites and glaciolacustrine varvites resepctively). Modern sedimentological investigations argue that diamictites are debrites formed by submarine slumping and argillites are turbidites with no paleoclimatic significance. Clustered and solitary Ediacaran microfossils of deep marine origin are present within argillites. Nonetheless, a glaciomarine setting correlative with a notionally global or near-global ‗Snowball Earth-type‘ glaciation, known as the Gaskiers glaciation (c.

580 Ma) is still favored by some workers based on a single exceptional outcrop of a thin (< 15 cm thick) succession of laminated pebbly argillites (< 2 cm thick laminae) comprising a very thin lower layer of matrix-supported diamictite containing supposed ice-rafted clasts (outsized

97 clasts8) and a conformable micro-thin upper massive mudstone layer. These facies are exposed at the type section of the Squantum Member at Squantum Head.

This paper presents detailed lithofacies data from the Squantum Head section and a depositional model that recognizes ‗ice-rafted dropstone horizons‘ as non-glacial ‗co-genetic debrite-turbidite couplets.‘ These form where thin dilute debris flows accelerate and evolve downslope into turbulent flows resulting in a couplet composed of a lowermost lamina of matrix-supported clasts (debrite) sharply or gradationally overlain by a micro-thin lamina of graded or massive argillite (turbidite). A review of the literature presented in this study shows that the global sedimentary record of ice rafting and floating ice during Gaskiers glaciation (c.

580 Ma) appears to be limited, inconsistent with the hypothesis of a global or near-global scale

‗Snowball Earth-type‘ glaciation that led to the growth of thick ice covers on the world‘s oceans.

8 Clasts that exceed the thickness of the enclosing bed or laminae. 98

3.1 INTRODUCTION AND PURPOSE OF STUDY

Neoproterozoic glaciations (c. 800-570 Ma) are considered by some to have been exceptional events unlike subsequent ice ages. These glaciations are interpreted to have been globally widespread extending from high to low latitudes with frozen oceans with estimated sea ice thicknesses of up to 400 m and a hydrosphere that was entirely shut-down (Hoffman et al., 1998; Evans, 2000; Hoffman & Schrag, 2000; 2002; Kirschvink, 2002; Schrag et al., 2002;

Hoffman, 2005; Hoffman et al. 2006; ‗Snowball Earth Hypothesis‘). One such cooling event

(out of a total of four postulated ‗Snowball Earths‘ between 770 and 570 Ma; Hoffman et al.,

2006; Stern et al., 2006; MacDonald et al., 2010) is referred to as the ‗Gaskiers glaciation‘ and dated at c. 580 Ma (Bowring et al., 2003; MacGabhann, 2005). It is named after the undisputed glaciomarine deposits of the Gaskiers Formation preserved within the Avalon Terrane of eastern Newfoundland, Canada. The Gaskiers Formation is dominated by turbidites and debris flows deposited in a deep water rift basin in a volcanic arc setting (Williams & King, 1979;

Gardner & Hiscott 1988; Eyles and Eyles, 1989) and contains a clear signal of a glacial influence on sedimentation in the form of striated and faceted dropstones within the succession

(Bruckner and Anderson, 1971; Bruckner, 1977; Williams and King, 1979; Eyles and Eyles,

1989).

The Gaskiers Formation is classically correlated with other Late Proterozoic (c. 670-

550 Ma) diamictite-bearing successions preserved in Neoproterozoic rift basins within the

Avalonian-Cadomian Terranes, which developed while attached or adjacent to the active northern margin of Gondwana at c. 620-570 Ma. Such formations include the Konnarock

Formation of southwestern Virginia, USA (Miller, 1994) (formally the upper Mount Rogers

Formation; Rankin, 1993), the Grandfather Mountain Formation of the southern Appalachians in eastern USA (Schwab, 1981; O'Brien et al., 1983; Rast & Skehan, 1983), the Granville

Formation of northern France (Winterer, 1963; Schwab, 1981) and the Squantum Member of the Boston Basin, Massachusetts, USA (Billings, 1976; Rehmer and Roy, 1976). A glacial interpretation for all these successions has long been questioned (Dott, 1961; Winterer, 1963;

Caldwell, 1964; Schermerhorn, 1974; Schwab, 1981; Eyles and Eyles, 1989; Eyles, 1990;

Bailey and Bland, 2000). In the case of the Squantum Member, which is a thick diamictite deposit (up to 218 m thick in the subsurface), recent work confirms a non-glacial origin for the diamictites (Chapter 2 of this thesis); however, there still remains the question of so-called

‗ice-rafted dropstones‘ reported by Sayles and LaForge (1910) and Sayles (1914, 1919) within laminated pebbly argillite units found in association with the Squantum Member. The uncritical acceptance of the glacial interpretation of these laminated pebbly argillites by some subsequent workers have resulted in these rocks being used as compelling evidence of a glaciomarine setting for the Boston Basin during the Neoproterozoic (Socci and Smith, 1987, p.36; Passchier and Erukanure, 2010). Some workers even propose that the Squantum Member belongs to a severe Gaskiers glaciation of worldwide extent (Thompson and Bowring, 2000;

MacGabhann, 2005; Stern et al., 2006; Kawai et al., 2008; Hoffman and Li, 2009; MacDonald et al., 2010) (see Carto and Eyles, 2011 for review), though the glacial interpretation of these rocks has been strongly of disputed by others (e.g., Dott, 1961; Bailey, 1984; Bailey and Bland,

2000).

The objective here is to critically re-evaluate the controversial argillite facies of the

Boston Basin in the light of a modern understanding of sedimentary processes and facies in cold ice-infested waters, and models for the emplacement of outsized clasts in deep marine

100 mass flow dominated successions. The broader significance of the results for Neoproterozoic paleoclimate models is then discussed.

3.2 GEOLOGICAL SETTING, STRATIGRAPHY AND AGE OF THE BOSTON BASIN

The Boston Basin, located in eastern Massachusetts (Fig. 3.1), is a complex, fault- bounded Neoproterozoic volcanic arc-type (back- or intra-arc) basin (Cardoza et al., 1990;

Thompson, 1993; Thompson et al., 1996) that formed within the Avalon Terrane while it, along with several other terranes, was attached or adjacent to the mid-latitude northern margin of Gondwana (c. 620-570 Ma). These terranes collectively formed the Avalonian-Cadomian

Orogenic Belt (Nance et al. 1991; Thompson et al., 2007). Following rifting of the terranes from Gondwana, the Avalon Terrane accreted to eastern Laurentia around c. 440 Ma (Trench and Torsvik, 1992; Dalziel et al. 1994; Cocks, 2000) and now underlies much of eastern USA.

The c. 650 to 600 Ma Dedham and Westwood Granitoids and felsic volcanic rocks of the

Lynn-Mattapan Volcanic Complex 596 ± 2 Ma (Rast and Skehan, 1983; Durfee Cardoza,

1990; Thompson et al., 1996) comprise the basement rocks of the Boston Basin (Pe-Piper and

Murphy, 1989; Smith and Socci, 1990; O‘Brien et al., 1996; Murphy et al., 1999). These volcanic units are overlain by the younger mafic to intermediate flows, pillows, and pyroclastics of the Brighton Volcanics (c. < 587 Ma; Thompson and Grunow, 2004) which are interbedded with the deep marine sedimentary facies of the Boston Bay Group in a series of east-northeast-trending graben sub-basins (Kaye & Zartman, 1980; Zartman & Naylor, 1984;

Thompson, 1993).

101

The marine volcanic-sedimentary fill (~ 7 km) of the Boston Basin consists of thick mass flow conglomerates and diamictites, sandstones and argillite turbidites with interbeds of volcanic flows and volcaniclastic debris (subaqueously reworked lapilli tuff beds) indicating that clastic sedimentation occurred coevally with volcanic activity along the Avalon Terrane

(Dott, 1961; Thompson, 1993; Chapter 2). The characteristics of these facies are summarized in Table 3.1 and illustrated in Figures 3.2A-L.

Sedimentological data indicates that the basin fill is derived from local sediment sources (e.g., Dedham and Westwood Granitoids, felsic volcanics of the Lynn-Mattapan

Volcanic Complex, and Westboro Formation of the surrounding uplands) (Dott, 1961; Billings,

1976; Thompson, 1993; Chapter 2 of this thesis). Diamictites (debrites) are represented by thick (up to 215 m) beds of homogeneous diamictite and heterogeneous (chaotic) diamictites recording the initial mixing of gravel and mud to create debrites. Diamictites are composed of very poorly sorted mixtures of subangular to subrounded volcanic and intrabasinal clasts (5-60 cm in diameter but with occasional boulders) supported by a large volume of poorly sorted matrix of clay-to sand-sized detritus (> 80 volume % of rock). Clast types and lithologies in conglomerates and diamictites are identical and typical of a local volcanogenic source area.

Both diamictite facies types are conformably interbedded with sandstone and argillite turbidites and characterized by common slump and soft-sediment deformation structures typical of a tectonically active slope setting. Associated argillites are composed by reworked volcanic ash and contain clustered and solitary Ediacaran microfossils (Bavlinella sp., Vidal, 1976; Lenk et al., 1982; Thompson and Bowring, 2000) and Aspidella terranoviva (Bailey and Bland, 2000).

Lahee (1914) and co-workers (Sayles and LaForge, 1910; Sayles, 1914, 1919) reported ‗rare‘ and ‗faintly‘ striated clasts from the Squantum Member but all later workers, including the

102 author, have failed to identify the presence of a single convincing striated or glacially shaped clast (Dott, 1961; Bailey et al., 1976; Bailey, 1984; Chapter 2 of this thesis).

The type section of the Squantum Member is in Quincy, Massachusetts, USA, where a small but relatively continuous outcrop is exposed along some 100 m of the shoreline at

Squantum Head (Fig. 3.3). The outcrop exposes interbedded diamictites, conglomerates and sandstones interbedded with intervals of finely laminated (< 1 mm thick) and normally graded argillites and laminated (< 2 cm thick) pebbly argillites. The thin intervals (composite thickness of < 15 cm) of laminated pebbly argillite contain scattered outsized clasts and, as a result, have long been interpreted as an ice-rafted facies due to their ‗varve-like‘ appearance

(Sayles, 1914, 1919; Wolfe, 1976; Cameron, 1979; Socci and Smith, 1987; Socci and Smith,

1990) despite arguments made by other workers for a sediment-gravity flow origin (Dott,

1961; Bailey et al., 1976; Bailey, 1984; Bailey and Bland, 2000).

3.3 FIELD AND LABORATORY METHODS

This study emphasizes the broader sedimentary context of ‗ice rafted facies‘ exposed at the classic section at Squantum Head (Fig. 3.3). Fieldwork involved employing standard sedimentary logging procedures (Miall, 1978; Eyles et al., 1983; Nicholas, 1999), which involved documenting clast size, shape and texture, the geometry/orientation of beds and bedding contacts, as well as the vertical/ lateral facies inter-relationships with overlying and underlying facies. Further insights were gained through detailed examination of cut slabs and thin-section analysis of small-scale sedimentary characteristics such as clast lithology, shape and orientation, and matrix petrology. A statistical analysis of the relationship between laminae

103 thickness and maximum clast size for the pebbly argillite layers was also conducted. Studies show that both more and coarser detritus can be carried in suspension in a flow of greater turbulence than in a flow of lesser turbulence (e.g., Dott, 1963; Potter and Scheidegger, 1966;

Nemec and Steel, 1984; Martins-Neto, 1996). As a consequence, deposits of sediment-gravity flows display a significant positive linear correlation between their laminae (or bed) thickness and the maximum clast size within the laminae or bed (Gloppen and Steel, 1981; Nemec and

Steel, 1984; Martins-Neto, 1996).

3.4 DESCRIPTION OF LAMINATED ARGILLITE FACIES AT THE SQUANTUM HEAD SECTION

The Squantum Head section exposes laminated argillite, laminated pebbly argillite, graded and massive sandstone, massive conglomerate, and massive (homogeneous) and crudely mixed (heterogeneous) diamictite facies. Strata strike NE and dip 45o SW.

Syndepositional slump, fold and load features occur in all units consistent with a subaqueous mass flow origin as discussed in Chapter 2 of this thesis. Conglomerates are debrites and sandstones and argillites are turbidites (e.g., Dott, 1961; Socci and Smith, 1989; Chapter 2 of this thesis). Homogeneous massive diamictite facies are debrites resulting from the mixing of gravel and mud during transport downslope; crudely mixed ‗heterogeneous‘ facies show incomplete mixing of clasts and mud, rip-up clasts and slabs of folded argillite and slump and load features (Crowell, 1957; Dott, 1961; Bailey et al., 1976; Socci and Smith, 1987; Chapter 2 of this thesis).

At the Squantum Head section a thin succession (cumulative thickness of 52 cm) of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies

104

(Figs. 3.4 A,B) conformably interbed a succession of massive diamictite. Laminated argillites are composed of normally graded muddy siltstone laminae typically < 1 mm thick (Figs. 3.4B,

D, E, 3.5A). Most laminae are planar and parallel but some are wavy and convoluted; starved ripples and loaded and slightly scoured and erosional bases are common. Laminations are composed petrographically of subangular to subrounded, clay-to silt-sized quartz and feldspar grains, fine iron oxide grains, and rare zircons. Very fine granules (< 0.1 mm in diameter) occur within some laminae but they are extremely scarce and rarely exceed the thickness of the enclosing lamina (Figs. 3.4E and Figs. 3.5A, B). These laminated facies occur together in composite units (3-8 cm thick) and are similar to other laminated argillite facies that dominate the fill of the Boston Basin which are interpreted as turbidites and have been mapped as the

‗Cambridge Argillite‘ (Dott, 1961; Thompson and Bowring, 2000). Overall, these facies are thickest in the northern region of the basin consistent with paleocurrent data that indicate an overall transport direction to the north into deep water (Socci and Smith, 1987; Lindsay et al.,

1970).

Pebbly argillite layers at Squantum Head are of very restricted thickness ranging from

0.5 cm to 2 cm thick, bundled together into composite units from 8 to 15 cm thick (Figs. 3.4B,

D, F, G and 3.5A). These units are conformably interbedded with the normally graded, laminated argillite units described above. Rarely, a single pebbly argillite lamina occurs within units of laminated, normally graded argillite (Fig. 3.4E and Fig 3.5A). Notably, each pebbly argillite layer forms a distinct ‗couplet‘ composed of a very thin lower layer of diamictite overlain by a microlaminated massive mudstone lamina (e.g., Figs. 3.4F, G and Figs. 3.5A, C).

The thickness of the uppermost mudstone lamina is variable; in some cases it is even and continuous for the entire length of the underlying diamictite lamina but it is also common for

105 this layer to be ‗patchy‘ and laterally discontinuous. Clasts comprise up to 25 % of the volume of any one diamictite lamina (basement derived granitic, basalt, rhyolitic, and andesitic clasts, and intrabasinal laminated argillite and massive sandstone clasts) supported by a cryptocrystalline, sericitic-rich purple-brown coloured matrix (75 vol. % rock).

Petrographically and texturally, the matrix material of both the upper and lower layers is identical to that within the normally graded, laminated argillites found above and below with the only exception being the higher sand content of the matrix within the pebbly argillite layers. A low number (< 20 %) of clasts show a strong preferred long axis orientation parallel or sub-parallel to lamina bounding surfaces (strike NE/ dip 45oSW) (Figs. 3.6A, B). Clast size within the lower diamictite lamina is restricted, ranging from granule- to pebble-sized detritus with the majority of the clasts being granule-sized (< 0.5 mm in diameter). Clasts are subrounded to rounded in shape (Figs. 3.5A-D) with the largest clasts (< 3 cm) being well rounded.

Clasts either ‗float‘ within the matrix or are concentrated toward the base of pebbly argillite laminae where clasts occur in contact with each other; they are conformable within individual lamina. Seldomly, individual lamina are depressed9 below larger clasts with succeeding lamina draped over the clast creating a distinct ‗pinch and swell‘ appearance (Fig.

3.5D and Figs.3.7A-C).

9 Depression takes the form of a simple downfold, conformable with the base of the clast. 106

3.5 INTERPRETATION OF FACIES AND RECONSTRUCTION OF

DEPOSITIONAL SETTING

As previously argued by Dott (1961) and Bailey et al. (1984), laminated and normally graded argillite facies are the product of low-density turbidites. The primary signature of turbidity current deposition is the presence of graded bedding (e.g., Prior and Bornhold, 1990;

Eyles and Eyles, 2000; Lowe and Guy, 2000; Mulder and Alexander, 2001). As discussed above such facies dominate the fill of the Boston Basin (e.g., ‗Cambridge Argillite‘) and have a total cumulative thickness of 5.5 km out of a total basin fill of ~7 km. The presence of rare and isolated clasts within turbidites, like those present in the laminated, normally graded argillites at Squantum Head (rare granules), are well reported (e.g., Hein, 1982; Miall, 1985; Bennet et al., 1996; Eyles and Eyles, 2000; Zaleha and Wiesemann, 2005; Eyles and Januszczak, 2007;

Oard, 2008). Such facies are the product of clasts being moved aloft and within the main body of the turbidity current (‗lagstones‘; Postma et al., 1988; Kim et al., 1995; Johansson and Stow,

1995; Crowell, 1999) or clasts abandoned (‗stranded stones‘) by flows continuing to move downslope and progressively thinning and fining (Donovan and Pickerill, 1997) or moving in front of the flow (‗outrunner clasts‘; Prior et al., 1982; Bornhold and Prior, 1990).

In the case of the laminated pebbly argillite ‗couplet‘ facies at Squantum Head, a key observation is that the lowermost diamictite laminae of these layers are no different lithologically and compositionally from their thicker diamictite counterparts that occur stratigraphically above and below in outcrop (e.g. compare Fig. 3.2H and Fig. 3.5A). In fact, clast and matrix in both argillite facies types are identical and are from the same intrabasinal sedimentary and volcanic sources indicating a clear genetic relationship. Significantly, the

107 lowermost diamictite laminae of the pebbly argillite layers are composed of clasts of a very restricted size range and contain no clearly ‗outsized‘ clasts (e.g., boulders) that penetrate underlying laminae or striated and faceted clasts, which are features typical of ice-rafted facies

(Thomas and Connell, 1985).

As stated, the thicker diamictites beds that occur above and below the succession of laminated argillite facies are debrites produced by the slumping and downslope mixing of gravel and mud (Dott, 1961; Socci and Smith 1987; Chapter 2 of this thesis). Accordingly, the so-called ‗ice rafted‘ horizons (laminated pebbly argillite ‗couplet‘ units), which occur conformably with these thick diamictites, are readily explained as the product of thin, relatively dilute debris flows in which small clasts were suspended in a muddy viscous matrix. These layers are, in short, ‗intraformational debrites‘ deposited conformably within successions of turbidites (laminated, normally graded argillites). The concentration of clasts at the base of the lamina and those found in a ‗floating‘ position within the lamina are features commonly observed in sediment-gravity flows where some clasts tend to concentrate near the flow base while others remain in a suspended position supported by matrix strength (e.g., Govier and

Aziz, 1972; Johansson and Stow, 1995). Downward bending and arching of laminae around some clasts has formerly been regarded as strong evidence of an ice-rafted origin (Sayles and

LaForge, 1910; Cameron and Jeanne, 1976; Socci and Smith, 1990). However, such features are not uniquely diagnostic of ice-rafted debris; it has been noted that in some cases the matrix strength of dilute debris flows may be insufficient to support larger clasts which will sink to the base of the flow and subsequently penetrate into the underlying soft sediment (Donovan and

Pickerill, 1997). In this case, the deeper penetration of a few larger clasts can be simply attributed to density contrasts between the clasts and surrounding matrix (Gilbert, 1990;

108

Doublet and Garcia, 2004). It is also important to note that depression of laminae rarely extends below the base of the clast, whereas ice-rafted dropstones commonly depress the sediment below the base of the clast to a depth of more than half the vertical dimension of the clast (Thomas sand Connel, 1985).

As described above, each pebbly argillite lamina is composed of a lower diamictite lamina that is overlain by very thin massive mudstone lamina forming a distinct couplet. This is a characteristic of debris flows where dense, sediment-rich laminar debris flows entrain water and partially develop into turbulent flows. The result is a flow with coeval lobes, and the subsequent deposition of what has been called a ‗co-genetic debrite-turbidite couplets‘

(Schuppers and Martinius, 1994; Haughton et al., 2003; Sylvester and Lowe, 2004; Talling et al., 2004; Amy and Talling, 2006; Barker et al., 2008; Pyles and Jennette, 2009). Several mechanisms have been proposed for the development of turbulent flows from debris flows including liquefaction and an increase in the velocity of the debris flow leading to the onset of turbulence (Allen, 1971; Schwarz, 1982; Mohrig et al., 1999; Mulder and Alexander, 2001;

Sohn et al. 2002; Talling et al., 2004; Haughton et al, 2006). Strikingly similar facies have been generated by subaqueous flows of volcanogenic sediment which result in laminated matrix- supported diamictite layers overlain by thin laminae of fine-grained sediment (Fisher and

Schmincke, 1984, pp. 286-7) (Fig. 3.8). The alternations between laminated, normally graded argillites (turbidites) and laminated pebbly argillite facies (co-genetic debrite-turbidite couplets) likely reflects changes in the volume of sediment delivered to the basin. The discontinuity or ‗patchiness‘ of the overlying mudstone lamina within some pebbly argillite layers can be attributed to a discontinuity of the flow regime downcurrent or post-depositional erosion by succeeding thin debris flows (Bass, 2004).

109

A sediment-gravity flow origin for pebbly argillite layers is further indicated by a positive relationship between ‗maximum clast size‘ and ‗lamina thickness‘ in pebbly argillite layers (correlation coefficient of r =0.33) (Figs. 3.9A, B). Based on the methods of Sadler

(1982), laminae thickness was determined by averaging the minimum and maximum thickness of each lamina and the maximum clast size of each lamina was determined by measuring the longest axis of the largest clast present in each lamina. Maximum clast size was used because it is widely agreed that it best depicts flow competency (Rocheleau and Lajoie, 1974; Nemec et al., 1980; Steel and Thompson, 1983).

Data show much variability but these results, including the low correlation coefficient value, are comparable to other studies of clasts transported by debris flows (e.g., Dott, 1963;

Potter and Scheidegger, 1966; Lowe, 1982; Nemec and Steel, 1984). Variability (e.g., small grains in thick laminae) can be attributed to the limited size range of available clasts to match the competence of the flows (e.g., Potter and Scheidegger, 1966). In addition, scatter in the plot could also reflect deposition from ‗over-competent‘ flows that thinned leaving larger clasts stranded within a thinner flow, which is typical of non-cohesive flows (Nemec and Steel, 1984;

Donovan and Pickerill, 1997; Mulder and Alexander, 2001). Some laminae may also represent a composite amalgamated bed rather than a single episode of flow (Larsen and Steel, 1978). In light of these caveats, the positive relationship indicates a general correlation between bed thickness and maximum grain size that would not be expected where clasts had been ice-rafted.

Ice-rafted facies typically show a complete lack of correlation between the size of the clasts and the thickness of the laminae in which they lie (Thomas and Connell, 1985). Although detailed analysis is constrained by a lack of published quantitative data of ice-rafted deposits, the correlation between maximum clast size and laminae thickness for the pebbly argillites at

110

Squantum Head suggests that these coarse layers were deposited by a debris flow mechanism in which mixtures of fine and coarse sediments moved downslope. As such, each lamina results from a single flow event in which matrix strength, dispersion and turbulence were the main clast support mechanisms (Mulder and Alexander, 2001). Based on visual inspection,

Bailey (1984) also made note that the maximum clast sizes in the pebbly argillite layers at the

Squantum Head section are proportional to the thickness of the enclosing laminae and the data presented here supports his finding.

It can be suggested that laminated pebbly argillites record the downslope spillage of debris from the margin of a thicker debris flow lobes into areas of fine-grained sedimentation down or across slope. In this scenario, laminated pebbly argillites developed as thin 'dilute' debris flows that escaped or ‗out-distanced‘ larger flows. This process has been noted by

Nemec and Steel (1984) and Socci and Smith (1987) where subaqueously resedimented material is reworked from active slump scars or the unstable ‗noses‘ of debris flows farther upslope (‗gravity-winnowing‘) (Fig. 3.10). In this later case, the term ‗outrunner debrite‘ may be an appropriate term for the laminated pebbly argillites in the Boston Basin.

3.6 DISCUSSION

The laminated pebbly argillite facies outcropping at Squantum Head in the Boston Basin of eastern Massachusetts, USA, have long been regarded as being ‗ice rafted‘ (Sayles and

LaForge, 1910; Sayles, 1914, 1919). In contrast, the results of this study show that the laminated pebbly argillite facies are ‗co-genetic debrite-turbidite couplets‘ as reported from non-glacial settings, and the associated laminated, normally graded argillite facies are

111 turbidites. Other work in the Boston Basin has shown that the conglomerate, diamictite, sandstone and other argillite facies of the basin fill are the product of non-glacial mass flow

(Dott, 1961, p. 1302; Chapter 2 of this thesis). In this regard, it is worth reviewing other data that have been used to infer a glaciomarine setting for the Boston Basin. These include an outsized pebble and cobble identified within a succession of alternating sandstone and conglomerate beds in the Newton area by Wolfe (1976), Cameron and Jeanne (1976) and Socci and Smith (1990). Given that these two isolated clasts occurred in a sandstone bed, these authors hesitated to interpret them as ice-rafted dropstones (e.g., see Cameron and Jeanne,

1976; pg. 133). Furthermore, Bailey (1984) emphasized that these two clasts occurred on the same bedding plane immediately below conglomerate filled channels, and reinterpreted them as lag deposits. Intrabasinal rip-up clasts and blocks of laminated argillites at Mystic River in

Somerville, MA, were suspected by Rehmer and Roy (1976) as ice-rafted dropstones but later workers agreed that they record the erosion by vigorous turbidity currents (Bailey et al., 1976;

Bailey, 1984). In a similar case, a single granitoid lonestone described as floating in a layer of silt and fine-grained sand with an intrabasinal laminated argillite ‗rip-up‘ clast was reported by

Williams (2008) in exposures of the ―Cambridge Argillite‖ in the Hewitt‘s Cove area.

However, McMenamin and Beuthin (2008) stressed that its association with an intrabasinal rip- up clast was suggestive of sediment-gravity origin. He also emphasized that no other isolated and/or clustered clasts have been found in any other outcrops of Cambridge Argillite (see Dott,

1961; Bailey et al., 1976; Bailey, 1984; Passchier and Erukanure, 2010). In view of the evidence presented above, a cold climate, glacially influenced marine setting can be ruled out for the Boston Basin.

112

Evidence of ice covers around 580 Ma along the Avalonian-Cadomian Terranes is found in Newfoundland in the form of ice-rafted striated and faceted clasts within the Gaskiers

Formation but the age and global correlation of other deposits is weakly constrained and their tillite interpretations remain suspect (see Hambrey and Harland, 1981; Eyles and Eyles, 1989;

Williams and Schmidt, 2000; Condon et al., 2002; Eyles and Januszczak, 2004, 2007; Miller et al., 2007; Eyles et al., 2007; Eyles, 2008; Allen and Etienne, 2008; Kawai et al., 2008; Zhao et al., 2009; Zhao and Zheng, 2010). Certainly the lack of evidence for a major glacial influence on sedimentation within the Boston Basin and other arc-related diamictites-bearing successions within the Avalonian-Cadomian Terranes rules out a severe global Snowball Earth glaciation at this time.

Compilation of the sedimentary record of floating ice from Gaskiers time indicates only a restricted occurrence of ice-rafted deposits worldwide (Table 3.3; Fig. 3.11). This finding is significant because many models of Snowball Earth climates stress the fundamentally necessary requirement for kilometre-thick, drifting ice covers on the world‘s oceans (known as

‗dynamic ice‘ or ‗sea glaciers‘; e.g., Goodman and Pierrehumbert, 2003; Poulsen, 2003; Lewis et al., 2003; Pollard and Kasting, 2005; Fairchild and Kennedy, 2007). According to these models, marine ice covers increase planetary albedo providing a hypothetical trigger for hyper- arid severely and hyper-cold global glaciations. Either the geological record has not been examined in sufficient detail or the global reach of floating ice is greatly exaggerated by these models.

In ice-infested glaciomarine environments such as those well-known from Late

Paleozoic and Pleistocene deposits and modern glacially influenced marine environments, ice- rafted facies such as ‗rain-out diamictite‘ that generally occurs in isolation within fine-grained

113 sediment (blanket-like deposits of pebbly mudstone that accumulate passively on the sea floor from ice bergs and meltwater plumes), laterally extensive horizons of far-traveled outsized

‗dropstone‘ clasts (coarse gravel to boulders) dumped by icebergs or pack ice are common

(Domack, 1990; Moncrieff and Hambrey, 1990; Dowdeswell et al., 2000; Hambrey and

McKelvey, 2000). However, such features have not been observed in the Boston Bay Group.

An alternative argument that ‗Snowball‘ ice masses left no sedimentary record as a consequence of fully frozen oceans and the inability of ice to drift (Halverson et al., 2004) seems unlikely. It seems reasonable to assume that such facies and structures would be dispersed world-wide during each Snowball Earth event, especially during the early and late stages of such events when the mobility of floating ice would be maximized. We could also anticipate deposits akin to Heinrich Event layers left on the ocean floors by Late Pleistocene armadas of icebergs traveling several thousand kilometers from their source ice sheets (e.g.,

Heinrich, 1988; Bond and Lotti, 1995; Andrews et al., 1998). It could be expected that large volumes of debris would likely have been trapped within thick, ocean-going ice masses of the

Neoproterozoic as ice froze in near-shore areas and along the grounding lines of continental ice sheets (see Hoffman, 2005, and discussion in Eyles and Januszczak, 2007). Similarly, channel- like ice scours and associated deformed and mixed sediment horizons (isolated troughs of coarse-grained diamictite within finer grained sediment and ice-berg plough marks) produced by the grounding and subsequent decay of floating icebergs on the sea floor (see Thomas and

Connell, 1985; Eden & Eyles, 2001; Eyles and Meulendyk, 2008) are rarely reported from the

Neoproterozoic marine basins (see Table 3.3 and reference therein). A lack of facies and structures created by floating ice in the Neoproterozoic is a major problem too, for alternative climate models involving more limited ice covers (e.g., ‗Slushball‘ and ‗Softball‘ Earth

114 models; Runnegar, 2000; Olcott et al., 2005; Peltier et al., 2007) or of regional ice covers restricted to tectonically-uplifted areas (Eyles and Januszczak, 2004; Eyles, 2008); all these models predict large areas of glacier ice in direct contact with ocean waters.

3.7 CONCLUSIONS

The laminated pebbly argillite facies and associated laminated argillites outcropping at

Squantum Head in the Boston Basin of eastern Massachusetts, USA, have been regarded as

‗ice rafted‘ facies and glaciolacustrine varvites since 1914. However, the results of this study show that no evidence of a glacial influence on the deposition of these facies can be recognized. Instead, the results of this study confirm the turbidite interpretation for the laminated argillites and demonstrate that laminated pebbly argillites are comparable to ‗co- genetic debrite-turbidite couplets‘ as reported from non-glacial marine settings. Classically, these rocks and associated strata of the Boston Basin have been viewed as recording the presence of ice covers in eastern North America during the Gaskiers glaciation, the fourth in a series of notional global or near-global glaciation events between 770 and 570 Ma but an entirely non-glacial deep marine setting is indicated for the fill of the Neoproterozoic Boston

Basin. In general, the global sedimentary record of ice-rafting and floating ice during the

Gaskiers glaciation (c. 580 Ma) appears to be rather limited and is inconsistent with the hypothesis that ice covers were of global extent. Data presented here demonstrates that a non- glacial interpretation for the laminated and normally graded argillites and laminated pebbly argillites at the Squantum Head section can not be ruled out and, therefore, bring into question the validity of the notion of significant ice development within the Boston Basin during the

115

Neoproterozoic. These results also bring into question the premise of paleoclimate models that stress the key role of drifting ice covers on the oceans (‗dynamic ice,‘ ‗sea glaciers‘) in changing planetary albedo and triggering a global glaciation at this time. The Gaskiers glaciation appears to have been restricted in area, not unlike Phanerozoic ice ages.

116

Figure 3.1 Location of the Boston Basin in eastern Massachusetts, USA (modified from Socci & Smith, 1987). The location of the study site at Squantum Head is indicated on map (Site 1) as well as the locations of lithofacies shown in Figures 3.2A-L. Site 1: Squantum Head; Site 2: Franklin Park; Site 3: Chestnut Hill Mall; Site 4: Webster conservation Area; Site 5: Arnold Arboretum; Site 6: Hewitts Cove. Local map (satellite image) of Squantum Head shown in Figure 3.3.

117

Figure 3.2 (A) Massive, clast-supported cobble-to-boulder conglomerate (Site 2; Fig. 3.1). (B) Massive, matrix-supported pebble-cobble conglomerate supported by coarse sand matrix (Site 3; Fig. 3.1).

118

Figure 3.2 (C) Stratified pebble-cobble conglomerate with coarse-grained sandstone (Site 4; Fig. 3.1). (D) Succession of laminated and normally graded sandstones (Site 1; Fig. 3.1).

119

Figure 3.2 (E) A succession of normally graded sandstone laminae deformed by large overturn fold (Site 1; Fig. 3.1). (F) Massive fine-grained sandstone underlying (contact at base of hammer) heavily weathered diamictite bed (Site 1; Fig. 3.1).

120

Figure 3.2 (G) Medium-to fine-grained sandstone showing low-angle cross-lamination (Site 4; Fig. 3.1). (H) Matrix-supported, homogeneous diamictites (Site 5; Fig. 3.1).

121

Figure 3.2 (I) Clast-supported diamictite showing crude normal grading, passing upward into matrix-supported diamictite (Site 1; Fig. 3.1). (J) Heterogeneous, matrix-supported diamictite (Site 1; Fig. 3.1).

122

Figure 3.2 (K) Laminated and normally graded argillite (Site 6; Fig. 3.1). (L) Bed of reworked lapilli tuff within a succession of laminated and normally graded argillite (Site 1; Fig. 3.1).

123

Study site Quincy Bay

Squantum Head

Figure 3.3 Satellite image of Squantum Head in Quincy, MA showing location of study site (site 1; Fig. 3.1) (from Google, Inc. 2010).

124

Figure 3.4 (A) Study site at Squantum Head where a succession of matrix-supported diamictite is interbedded with a thin succession (composite thickness of 52 cm) of alternating composite units of laminated pebbly argillite and laminated, normally graded argillite (location outlined in dashed rectangle). (B) Enlargement of a sequence of alternating units of laminated, normally graded argillite (top) and laminated pebbly argillite (below).

125

raphic log of a a of log raphic

F e mudstone. mudstone. e

og of study section shown in Figure 3.4C. 3.4C. in Figure shown section study of og

laminae of massiv of laminae

p of study site showing two massive diamictite beds interbedded by the thin succession of alternating units of laminated pebb oflaminated units alternating of succession thin the by interbedded beds diamictite massive two showing site study of p

composite unit of laminated pebbly argillite shown in Figure 3.4D. (G) Diagram of pebbly argillite layers (shown in Figure in Figure (shown layers argillite of pebbly (G)Diagram 3.4D. inFigure shown argillite pebbly laminated of unit composite

argillite and laminated and normally graded argillite (heavily eroded area). (D) Stratigraphic l (D)Stratigraphic area). eroded (heavily argillite graded normally and laminated and argillite

(E) Stratigraphic log of a single composite unit of laminated and normally graded argillite shown in Figure 3.4D. (F) Stratig (F) 3.4D. inFigure shown argillite graded normally and laminated of unit composite single a of log (E)Stratigraphic single micro upper an and ofdiamictite laminae lower of a composed unit each showing 3.4F)

Figure 3.4 Figure 126

Figure 3.5 (A) Cut slab of alternating units of laminated, normally graded argillite and laminated pebbly argillite facies at Squantum Head. Length of slab is 20 cm. Arrow indicates ‗way-up‘ direction.

127

Argillite with granules

Argillite

Scale |:------| 400 Microns

Silty-mudstone

Diamictite

Scale |:------| 400 Microns

Figure 3.5 (B) Photomicrograph of a thin section of normally graded argillite laminae showing contact between successive lamina (dashed line) (polarizers uncrossed). ‗Way-up‘ direction is indicated by black arrowhead. (C) Photomicrograph of a thin section of a single pebbly argillite lamina showing contact with overlying silty mudstone lamina (polarizers uncrossed).

128

I II

III

Figure 3.5 (D)Photograph of a cut-slab of laminated pebbly argillites indicating large igneous clasts that depress the underlying laminae or are ‗enveloped‘ by underlying and overlying laminae (e.g., clasts I, II, and III). Note: clasts within individual lamina are typically in a ‗floating‘ position or occur along the basal surface of the laminae. Slab segment is 6 cm in length and is 7 cm in width.

129

Laminated and normally graded argillite

Laminated

pebbly argillite B

Lower diamictite bed

B I

Figure 3.6 (A) Succession of laminated, normally graded argillite and laminated pebbly argillite overlying a bed of diamictite at study site (see Figure 3.4A). (B) Close-up of ‗outsized clasts‘ (I and II) within laminated pebbly argillites (shown in Figure 3.6A) with long axes parallel to basal surfaces of lamina (strike NE/dip 45o SW).

130

A B

Scale |:------| 400 Microns Scale |:------| 400 Microns

Scale |:------| 400 Microns

Figure 3.7 (A) Photomicrograph of an argillaceous clast within the lower diamictite lamina of a pebbly argillite layer where no clear depression of sediment around the clast is present. (B) Photomicrograph of quartzose clast within the lower diamictite lamina of a pebbly argillite layer creating a distinct ‗pinch and swell‘ appearance with overlying mudstone lamina. Note: laminae below this clast are not depressed. (C) Photomicrograph of a lithic clast within the lower diamictite lamina of a pebbly argillite layer creating a distinct ‗pinch and swell‘ appearance at the upper boundary of the enclosing laminae. ‗Way-up‘ direction is indicated by black arrowheads. 131

Figure 3.8 Thinly-bedded, subaqueous Devonian laminae of basaltic composition deposited during submarine eruptions (Madfeld, Germany) (from Fisher and Schmincke, 1984).

132

1.60

1.40 R2 = 0.3344 1.20

1.00

0.80

0.60

Maximum size clast (cm) 0.40

0.20

0.00 0.00 0.20 0.40 0.60 0.80 1.00 1.20 1.40 1.60 1.80 Laminae thickness (cm)

Figure 3.9 (A) Composite unit of pebbly argillite layers measured for ‗Maximum clast size versus laminae thickness‘ analysis. Boundaries of each lamina (drawn in black lines) and clasts used for measurement (lettered and lined) are indicated. Refer to Table 3.2 for clast size and bed thickness data. (B) Scatter-plot of ‗Maximum clast size and lamina thickness‘ data for the composite sequence of laminated pebbly argillites shown in Figure 3.9A, showing ‗best-fit‘ linear regression line and correlation coefficient.

133

Unknown water depth

Figure 3.10 Conceptual model of the evolution of ‗co-genetic debrite-turbidite‘ layers (laminated pebbly argillites) exposed at Squantum Head, as described in the text. Diagram shows how laminated pebbly argillites developed as thin, relatively dilute debris flows in which small clasts were suspended in a muddy viscous matrix due to the downslope spillage of debris that escaped or ‗out-distanced‘ active slump scars or the unstable ‗noses‘ of thicker debris flows farther upslope.

134

Figure 3.11 (A) Paleogeographic map of microcontinents at c. 580 Ma showing the distribution of Gaskier-aged glacigenic formations (stars) identified by Hoffman and Li (2009) (modified by Li et al., 2008). (B) Reconstruction of microcontinents ~565 Ma showing locations of Gaskiers glacial deposits reported by Kawai et al. (2008). Eg: Egan Formation (Corkeron and George, 2001), Et: Elatina Formation (Calver et al., 2004), Gs: Gaskiers Formation (Bowring et al., 2003), Hk: Hankalchough (Xiao et al., 2004), JG: Jibalah Group (Stern et al., 2006), LC: Loch na Cille Boulder Bed (Condon and Prave, 2000), Mt: Mortensnes Formation (Halverson et al., 2005), Mv: Moelv Formation (Bingen et al., 2005), SA: Serra Azul Formation (Alvarenga et al., 2007), Sq: Squantum Member (Thompson and Bowring, 2000), ANS; Arabian-Nybian Shield (Stern et al., 2006).

135

Table 3.1: Descriptions and interpretations of lithofacies of the Boston Bay Group.

136

Table 3.1: (continued)

137

Table 3.2: Maximum clast size (diameter of longest axis) and laminae thickness data collected from laminated pebbly argillite facies (Fig. 3.10). Note: Bed number 13 was omitted due to laminae distortion.

138

Table 3.3: Summary of sedimentary formations widely assigned to the Gaskiers glaciation at c. 580 Ma (based on those reported by Hoffman and Li, 2009 and Kawai et al., 2008). Formations where possible dropstones have been reported are highlighted in grey.

Description "Suspected" Gaskiers depositet (Kawai al. 2008). by mudstone- laminated diamictiteMassive m) overlain (70 siltstone m) which contain (200 sparse granules. "Suspected" Gaskiers depositet (Kawai al.2008). diamictiteMassive m) consisting(24.6 sandstone, of pebbly brecciated of sandstone interbeds with conglomerates and minor limestone flaggy dolomite laminated interbeds. and and cobble-sized in 'dropstones" and dolomite Pebble beds. in sand, silt mud facies. or sandstones into thatsiltstonelaminated Fluvial grade lonestones, mudstone with red rare and and pass laterally into diamictite ( 5m). Slumped common. are soft-sediment features and Diamictite intercalations siltstone, mudstone laminated and with of rare aswith lonestones dropstones. identified Diamictite, clastscrudely stratified with angular limonite or in carbonate matrix volcaniclastic and siltstonesandstone. Thin sandstone lensesand cm (30 thick). clastsTill sandstone-siltstone in graded finely beds. as Deposit tillite,asinterpreted but re-interpreted has been area. glacially-influenced from deposit derived flow debris diamictite massive Thick beds of by separated sandstone units beds, thinly thinly-bedded of argillites, siltstones,bedded siltstones tuffaceous mudstonesand containing small isolated ("dropstones"). pebbles Volcanicbombs, agglomerates, air-fall blocks lapilli within succession.and ash identified diamictiteby Massive sandstone,conglomerate, pebbly siltstoneinterbedded rhythmicallyand argillite laminated facies. Outsized clasts in pebbly siltstones. identified diamictitesCoarse to conglomerates (1 or m),40 siltstones,laminated finely sandstonesand with scattered pebbles. Most diamictites mixitites. are within diamictites. observed been outsizedor Dropstones clastsnever have siltstonesLaminated as distalinterpreted are turbidites.and Slumped features turbiditessuccession of sediment of downslope. suggest reworking diamictiteMassive m) (<50 with laminatesmudstones. of Formation. Stappogiedde of Member deposits into Lillevatan Grades non-glacial of environments shelf and by nearshore to deposits energy low high- of Overlain associatedLonetones with turbidites. sedimentary diamictiteand clasts Massive m) containing volcanic (5-15 by mudstone laminated laterallywith outsized extensive local of Overlain origin. clasts to up cm30 facies"). in ("dropstone diameter diamictitemassive of divisions to (7 m69 Large thick), sandstones, thin rhythmical clay alterations of siltstoneand with outsized clasts that "squeeze Lonestones within 1984)). cut and beds" the Chumakov, ("dropstones" underlying of sandstone interbeds. Massive, thick diamictites to (50 m80 thick) that occur with lithologically similar to conglomerates (8 10m thick) thinlyand siltstoneslaminated shales and to(0.5 4 m thick) that contain lonestones.as Laminates possibleinterpreted varvites. debris. as ice-rafted interpreted horizons' 'pebbly or beds' Discontinuous 'boulder Diamictitebyflow. sediment reworked gravity Striatedbasement, polished and as clasts.aswell striated, iron-shaped and faceted Rhythmites to with (up dropstones 7 cm). pyrite-rich diamictitesand Thick organic with striated clasts abundant and "dropstones" in diamictite. "Suspected Gaskiers deposit" et (Kawai al., Diamictite 2008). polymicticor m (10 thick),conglomerate occurs with sandstone,conglomerate, limestones, shale, and some with stromatolitic mats lenses. pebbly conglomerate and "Suspected Gaskiers et (Kawai diamictiteal,No 2008). deposits present, linked to been Gaskiers. onlyhave lithostratigraphic boundaries sequence - 250 m 250 m 400 m 150 m 500 <50 m m 400 m 300 8-215 m 8-215 m 10-60 m 4-200 50-83 m 50-83 30-200 m 30-200 250-300m 40 -100 m -100 40 m -270 70 250-300 m 250-300 Thickness - - - - - Group Togari Group/ Togari Charod Group Charod Group Jibalah Grassy Group Hedmark Group Hedmark Quruqtagh Group Quruqtagh Conception Group Conception Yerelina Subgroup Yerelina Verstertana Group Boston Bay Group Louisa Downs Group Louisa Downs Kanunnah Subgroup Kanunnah (Dalradian Supergroup) (Dalradian Southern Highland Group Highland Southern (USA) Shield Arabia Russia (Brazil) Boston (France) (Russia) Southern Australia (Canada) Normady Scotland Location Tasmania Tasmania (Australia) (Australia) North China North and Westernand Newfoundland Northern Urals Northern South South America Northern Norway Northern China Northwest Southern Norway Southern Northern Australia Northern Northeast Arabian Arabian Northeast Africa Northeast

Tarim Block Arabia Arabia Baltica Baltica Baltica Baltica Avalonia Avalonia Australia Australia Australia Australia Cadomia Laurentia Amazonia

China North Paleocontinent

Egan Moelv Shield Elatina Breccia Dhaiqa Cottons Granville Luoquan Vilchitsy

Gaskiers Squantum Serra AzulSerra Croles Hill Croles Mortensnes Loch na Loch Cille na Churochnaya Hankalchough Arabian-Nubian Arabian-Nubian Name of Formation 139

Table 3.3: (continued)

Reference

Bowring et al. 2003; Figueiredo et et al. al. Bowring Figueiredo Figueiredo 2004; 2003; et al., et 2007. al. Alvarenga 2006; 2001. George, and Cockeron Coats Preiss,1964; and Johnson, and Mawson, Dalgarno 1949; Gostin,and Lemon 1987; Eyles1990; et al., 2007. et al. 2004. Calver Walter,and Preiss, Calver 2000; et al. 2004. Calver 2000; King, and Anderson 1981; Anderson, Bruckner 1971; and Eyles Eyles,and et al.,Bowring 1989; 2003. 1981; Sayles Wolfe, 1976; Rehmer, 1976; Cameron, 1914; 2000. Bowring, and Thompson 1964. Winterer Graindor, 1963; Banks et al. Edwards, 1984; 1971; et al., 2001. Gorokhov Nystuen, Bjørlykke1976; et al., 1976. 1978. Chumakov, 1969; Chumakov, and Bessonova 1981. 1978, Chumakov, et al., Condon 1999; 2000. Prave, et Guan al., etShen 1986; al., 2007. et 1984; al., Zhu, and Zhao Gao 1980; et al., 2004. Xiao 2003 Corsetti Kaufman, and Stern et al., 2006. Stern et al., Miller 2006; et al., 2007.

Age 542Ma Vendian >640 >640 Ma <600 Ma <600 <600 Ma. Ma 582±4 Ma 575±3 Ma 580±1 584 ± 0.5 584 to Ma 650-600 Ma 595-570 Ma 640-580 to 615±6Ma Vendian age Vendian age Vendian 582 ± 0.4 582 Ma 660 to Ma660 670 Upper Riphean Upper (1100 to Ma) 680 (1100 (630 and 560 Ma) 560 and (630 Ma) 560 and (630 ~530 Ma and 600 Ma) 560 and ( 630

Belt) Belt)

Urals Basin Qinling Brittany Dalradian Kanunnah Kimberleys Boston Basin and Hedmark Basin Dhaiqa Alto Paraguay Gaissa Basin Valdres Valdres Basins

Smithton Basin Arc-related rift basinrift Arc-related Central Flinders zone Central Shield Arabian-Nybian (Riphean Orsha Basin)Orsha (Riphean Northern Paraguay belt Paraguay Northern Louisa Louisa Basin (syncline) Belt Orogenic Tianshan of Adelaide GeosynclineAdelaide of platform Eastern Europe in Azul the Serra syncline (upper brioverian geosyncline) brioverian (upper (Avalonian-Cadomian Orogenic (Avalonian-Cadomian Orogenic (Avalonian-Cadomian

Egan Moelv Shield Elatina Breccia Dhaiqa Cottons Granville Luoquan Vilchitsy

Gaskiers Squantum Serra AzulSerra Croles Hill Croles Mortensnes Loch na Loch Cille na Churochnaya Hankalchough Arabian-Nubian Name of Formation

140

REFERENCES

Allen, J.R.L. (1971). Mixing at turbidity current heads, and its geological implications. Journal of Sedimentary Petrology, 41, 97–113.

Allen, P. A. & Etienne, J.L. (2008). Sedimentary challenge to Snowball Earth. Nature Geoscience, 1, 817.

Alvarenga, C.J.S., Figueiredo, M.F., Babinski, M. & Pinho, F.E.C. (2007). Glacial diamictites of Serra Azul Formation (Ediacaran, Paraguay belt): evidence of the Gaskiers glacial event in Brazil. Journal of South American Earth Sciences, 23, 236–241.

Amy, L. A. & Talling P.J. (2006). Anatomy of turbidites and linked debrites based on long distance (120 x 30 km) bed correlation, Marnoso Arenacea Formation, Northern Apennines, Italy. Sedimentology, 53, 161-212.

Anderson, M.M. (1987) Late Precambrian glaciomarine sequence, Conception Group, Double Road Point, St Mary‘s Bay, southeast Newfoundland. In: Centennial Field 618 Nicholas Eyles and Carolyn H. Eyles Guide Vol. 5 (Ed.) D.C. Roy, Geologicial Society of America, 473- 476 pp.

Andrews, J. T., Kirby, M., Jennings, A. E. & Barber, D. C. (1998). Late Quaternary stratigraphy, chronology, and depositional processes on the slope of S. E. Baffin Island, detrital carbonate and Heinrich Events: Implications for onshore glacial history. G´eographie Physique et Quaternaire 52, 1– 15.

Ayres, L.D., Van Wagoner, N.A. & Ferreira, W.S. (1991). Voluminous shallow-water to emergent Phreatomagnetic Basaltic Volcaniclastic rocks, Proterozoic (~1886 MA) Amisk Lake Composite Volcano, in Flon Greenstone Belt, Canada. Sedimentation in Volcanic Settings, SEPM Special Publication, 45, 175-187.

Bailey, R.H., Newman, W. A. & Genes, A. (1976). Geology of the Squantum "Tillite": p. 92- 106, In: B. Cameron (Ed), Geology of Southern New England, 68th Annual Meeting New England Intercollegiate Geological Conference, Science Press, Princeton, NJ.

Bailey, R. H. (1984). Origin of Dropstones in the Precambrian Boston Bay Group. Geological Society of America Abstracts with Programs, 16(2) p. 2.

Bailey, R.H. & Bland, B.H. (2000). Ediacaran fossils from the Neoproterozoic Boston Bay Group, Boston area, Massachusetts. Geological Society of America Abstracts with Programs, 32, p. A11.

Banks, N.L, Hobday, D.K., Reading, H.G. & Taylor, P.N. (1974). Stratigraphy of the Late Precambrian `Older Sandstone Seriers‘ of Varangerfjord Area, Finnmark. Norges geologiske undersøkelse, 303, 1-15.

141

Bass, J.H. (2004). Conditions for formation of massive turbiditic sandstones by primary depositional processes. Sedimentary Geology, 166, 293–310.

Bennet, M.R., Doyle, P. & Mather, A.E. (1996). Dropstones: their origin and significance. Palaeogeography, Palaeoclimatology, Palaeoecology, 121, 331-339.

Bessonova, V. Ya. & Chumakov, N.M. (1968). On glacial deposits in the Late Precambrian of Belorussia. Dokl. Akad. Nauk SSR, 178(4), 68-71.

Billings, M.P. (1976). Geology of the Boston Basin. In: J.W. Skehan, & D.P Murray (Eds.), Studies in New England Geology. Memoir: Geological Society of America, 46, 5-30.

Bingen, B., Griffin, W.L., Torsvik, T.H. & Saeed, A. (2005). Timing of Late Neoproterozoic glaciation on Baltica constrained by detrital zircon in the Hedmark Group, south-east Norway. Terra Nova, 17, 250-258.

Bjørlykke, K., Elvsborg, A. & Høy, T. (1976). Late Precambrian sedimentation in the central sparagmite basin of south Norway. Norsk Geologisk Tidsskrift, 56, 233-290.

Bond, G.C. & Lotti, R. (1995). Iceberg discharges into the North Atlantic on millennial time scales during the last glaciation. Cience, 267, 1005-1010.

Bornhold, B.D. & Prior, D.B. (1990). Morphology and sedimentary processes on the subaqueous Noeick River delta, British Columbia, Canada. In: A. Colella & D.B. Prior (Ed.), Coarse-grained Deltas International Association of Sedimentologists Special Publications, 10, 169-181.

Bouma, A. H. (1962). Sedimentology of some flysch deposits, a graphic approach to facies interpretation, pp 168; Amsterdam: Elsevier.

Bowring, S.A., Myrow, P.M., Landing, E. & Ramezani, J. (2003). Geochronological constraints on terminal Neoproterozoic events and the rise of metazoans. NASA Astrobiology Institute General Meeting Abstracts, 113–114.

Bruckner, W.D. & Anderson, M.M. (1971). Late Precambrian glacial deposits in southeastern Newfoundland-a preliminary note. Geological Association of Canada Proceedings, 24, 95-102.

Bruckner, W.D. (1977). Significance of new tillite finds for east-west correlation of Proterozoic Avalon-zone formations in southeastern Newfoundland (Canada). Geology, 33, 357-364.

142

Caldwell, D.W. (1964). The Squantum Formation: Paleozoic tillite or tilloid? Guidebook to Field Trips in the Boston Area and Vicinity. In: J.W. Skehan (Eds.), New England Intercollegiate Geological Conference, 56th Annual Meeting, 53-60.

Calver, C.R. & Walter, M.R. (2000). The late Neoproterozoic Grassy Group of King Island, Tasmania: correlation and palaeogeographic significance. Precambrian Research 100, 299–312.

Calver, C.R., Black, L.P., Everard, L.P. & Seymour, D.B. (2004). U–Pb zircon age constraints on late Neoproterozoic glaciation in Tasmania. Geology, 32, 893–896.

Cameron, B. (1976). Geology of Southeastern New England, Field Trip Guidebook: New England Intercollegiate. Geological Conference, Boston University. Science Press, Princeton, N.J.

Cameron, B. & Jeanne, R.A. (1976). New Evidence for Glaciation During Deposition of the Boston Bay: p. 117-134. In: B. Cameron (Ed.), Geology of Southern New England, 68th Annual Meeting New England Intercollegiate Geological Conference, Science Press, Princeton, NJ.

Cardoza, K.D., Hepburn, J.C., & Hon, R. (1990). Geochemical constraints on the paleotectonic settings of two late Proterozoic mafic volcanic suites, Boston-Avalon Zone, Eastern Massachusetts, In: A.D. Socci, J.W., Skehan, and G.W. Smith (Eds.), Geology of the composite Avalon terrane of southern New England. Geological Society of America Special Paper, 245, 113-131.

Carto, S. L. & Eyles, N. (2011). The Squantum Member of the Boston Basin, Massachusetts, USA. In: E.Arnaud, G. P. Halverson, & G.A. Shields-Zhou (Eds.), The Geological Record of Neoproterozoic glaciations. Geological Society of London, Memoir, (in press).

Chumakov, N.M. (1978). Precambrian Tillites and Tilloids (in Russian). Nauka, Moscow.

Chumakov, N.M. (1981). Late Precambrian Kurgashlya tilloids, southern Urals, U.S.S.R. In: M.J. Hambrey, and W.B. Harland, (Eds.) Earth‘s Pre-Pleistocene Glacial Record. Cambridge University Press, Cambridge, p. 674-677.

Coats, R.P., Blissett, A.H., (1971). Regional and economic geology of the Mount Painter Province. Bulletin: Geological Survey of Southern Australia, 43, 426 pp.

Cocks, L.R.M. (2000). The Early Paleozoic geography of Europe. Journal of the Geological Society, 157, 1–10.

143

Condon, D.J., Prave, A.R. & Benn, D.I. (2002). Neoproterozoic glacial-rainout intervals: Observations and implications. Geology, 30, 1, 35–38.

Corkeron, M.L., & George, A.D. (2001). Glacial inclusion on a Neoproterozoic carbonate platform in the Kimberley region, Australia. Bulletin: Geological Society America, 113, 1121– 1132.

Corsetti, F.A., & Kaufman, A.J. (2003). Stratigraphic investigations of carbon isotope anomalies and Neoproterozoic ice ages in Death Valley, California. Bulletin: Geological Society of America, 115, 916–932.

Crowell, J.C. (1957). Origin of pebbly mudstones. Bulletin: Geological Society of America, 68, 993-1010.

Crowell, J.C. (1999). Pre-Mesozoic ice ages: their bearing on understanding the climate system. Memoir: Geological Society of America 192, 106 pp.

Dalgarno, R., Johnson, J.E., (1964). Glacials of the Marinoan Series, central Flinders Ranges. Geological Survey of Southern Australia Quaternary Geology, Notes 11, 3– 4.

Dalziel, I.W., Dalla Salda, L.H., & Gahagan, L.M. (1994). Paleozoic Laurentia-Gondwana interaction and the origin of the Appalachian-Andean mountain system. Bulletin: Geological Society of America, 106, 243–252.

Deynoux. M., Ciner, A., Monod, O., Karabiyikoglu, M., Manatschal, G. And Tuzcu, S. (2005). Sedimentary Geology, 173, 315–343

Domack, E. W. (1990). Laminated terrigenous sediments from the Antarctic Peninsula: the role of subglacial and marine processes. In: J. A. Dowdeswell, & J.D. Scourse (Eds.), Glaciomarine environments: processes and sediments. Geological Society of London, 91, 103.

Donovan, S.K. & Pickerill, R.K. (1997). Dropstones: their origin and significance: a comment. Palaeogeography, Palaeoclimatology, Palaeoecology, 131, 175-178.

Dott, R.H. Jr. (1961). Squantum ―tillite‖, Massachusetts—evidence of glaciation or subaqueous mass movements? Bulletin: Geological Society of America, 72, 1289-1305.

Dott, R.H. Jr. (1963). Dynamics of subaqueous gravity depositional processes. Bulletin: American Association of Petrological Geology, 47, 104– 128.

Doublet, S. & Garcia, J.P. (2004). The significance of dropstones in a tropical lacustrine setting, eastern Cameros Basin (Late Jurassic–Early Cretaceous, Spain). Sedimentary Geology, 163, 293–309.

144

Durfee Cardoza, K., Hepburn, J.C., & Hon, R. (1990). Geochemical constraints on the paleotectonic settings of two late Proterozoic mafic volcanic suites, Boston–Avalon zone, eastern Massachusetts. In: A.D. Socci, J.W. Skehan, & G.W. Smith (Eds.), Geology of the composite terrane of southern New England. Geological Society of America, Special Paper, 245, 113–131.

Eden, D. & Eyles, N. (2001). Description and numerical model of a Pleistocene iceberg scour at Scarborough Bluffs, Ontario, Canada. Sedimentology, 48, 1079–1022.

Edwards, M.B. (1975). Glacial retreat sedimentation in the Smalfjord Formation, late Precambrian, North Norway. Sedimentology, 22, 75–94.

Edwards, M.B. (1984). Sedimentology of the Upper Proterozoic Glacial Record, Vestertana Group, Finnmark, North Norway. Norges Geologiske Undersøkelse Bulletin, 394. 76 p.

Elles, G.L. (1934,). The Loch na Cille Boulder Bed and its place in the Highland succession. Geological Society of London, 91, 111–147.

Evans, A.D. (2000). Stratigraphic, geochronological, and palaeomagnetic constraints upon the Neoproterozoic climatic paradox. American Journal of Science, 300, 347–433.

Eyles, N. & Clark, B. M. (1985). Gravity-induced soft-sediment deformation in glaciomarine sequences of the Upper Proterozoic Port Askaig Formation, Scotland. Sedimentology, 32, 789- 814.

Eyles, N., & Eyles, C.H. (1989). Glacially influenced deep-marine sedimentation of the Late Precambrian Gaskiers Formation, Newfoundland, Canada. Sedimentology, 36, 601-620.

Eyles, N. (1990). Marine debris flows: Late Precambrian ‗‗tillites‘‘ of the Avalonian– Cadomian orogenic belt. Palaeogeography, Palaeoclimatology, Palaeoecology, 79, 73–98.

Eyles, C.H. & Eyles, N. (2000). Subaqueous mass flow origin for Lower Permian diamictites and associated facies of the Grant Group, Barbwire Terrace, Canning Basin, Western Australia. Sedimentology, 47, 343– 356.

Eyles, N. & Januszczak, N. (2004). ‗Zipper-rift‘: A tectonic model for Neoproterozoic glaciations during the break-up of Rodinia after 750 Ma. Earth Science Reviews, 65, 1-73.

145

Eyles, C.H., Eyles, N. and Grey, K. (2007) Palaeoclimate implications from deep drilling of Neoproterozoic strata in the Officer Basin and Adelaide Rift Complex of Australia: a marine record of wet-based glaciers. Palaeogeography, Palaeoecology, Palaeoclimatology 248, 291- 312

Eyles, N. & Januszczak, N. (2007). Syntectonic sub-aqueous mass flows of the Neoproterozoic Otavi Group, Namibia: where is the evidence for global glaciation? Basin Research, 19, 179– 198.

Eyles, N. (2008). Glacio-epochs and the supercontinent cycle after 3.0 Ga: Tectonic boundary conditions for glaciation. Palaeogeography, Palaeoclimatology, Palaeoecology, 258, 89–129.

Eyles, N. & Meulendyk, T. (2008) Ground penetrating radar study of a Pleistocene ice-scoured glaciolacustrine sequence boundary. Boreas, 37, 226-233.

Fairchild, I.J. & Kennedy, M.J. (2007). Neoproterozoic glaciation in the Earth System. Journal of the Geological Society, London, 164, 895-921.

Figueiredo, M.F, Babinski, M., Alvarenga, C.J.S. de, Pinho, F.E.C. (2004). Diamictites overlying Marinoan-age carbonates of Araras Formation, Paraguay belt, Brazil: evidence of a new glaciation?. In: Symposium on Neoproterozoic-Early Paleozoic Events in SW-Gondwana, 1, Extended Abstracts, IGCP Project 478, Second Meeting, Brazil, p. 18–19.

Figueiredo, M.F., Babinski, M., Alvarenga, C.J.S.de, Pinho, F.E.C., Simon, C.M. (2006). Chemostratigraphy (C, O and Sr) of Ediacaran post-glacial carbonates of Paraguay Belt, Mato Grosso, Brazil. In: South American Symposium on Isotope Geology, Uruguay.

Fischer, R.V., & Schminke, H.U. (1984). Pyroclastic Rocks. Springer-Verlag, Berlin.

Fisher, R.V. (1984) Submarine volcaniclastic rocks. In: B.P. Kokelaar, and M.F. Howells (Eds.), Marginal basin geology: volcanic and associated sedimentary and tectonic processes in modern and ancient marginal basins. Special Publications, Geological Society of London, 16, 5-27. Folk, R.L., (1974). Petrology of Sedimentary Rocks. Hemphill, Austin, TX, 182 pp.

Gao, Z., Zhu, S. (1984). Precambrian Geology in Xinjiang China. Xinjiang People‘s Publishing House, Urumuqi, China, 151 pp.

Gardner, S. & Hiscott, R.N. (1988). Deep-water facies and depositional setting of the lower Conception Group (Hadrynian), southern Avalon Peninsula, Newfoundland, Canadian Journal Earth Science, 25, 1579–1594.

Gilbert, R. (1990). Rafting in glacimarine environments. In: J.A. Dowdeswell & J.D. Scourse, (Eds.) Glacimarine Environments: Processes and Sediments. Geological Society of London Special Publications, 53, 105–120.

146

Gloppen, T.G., Steel, R.J. (1981). The deposits, internal structure and geometry in six alluvial fan-delta bodies (Devonian-Norway)- a study in the significance of bedding sequence in conglomerates. In: F.G. Ethridge y R.M. Flores (Eds.) Recent and ancient nonmarine depositional environments: Models for exploration. Society of Economic. Paleontologist and Mineralogists Special Publication, 31, 49-69.

Goodman, J. C. & Pierrehumbert, R. T. (2003). Glacial flow of Floating Marine Ice in `Snowball Earth‘. Journal of Geophysical Research, 108, C10, 3308.

Gorokhov, I.M., Melnikov, N.N., Negrutsa, V.Z., Turchenko, T.L. & Kutavin, E.P. (2001). Multistage illite evolution in Upper Proterozoic shales of Sredni Peninsula, the Murmansk coast of the Barents Sea. Lithology and Mineral Resources, 37,162-169.

Govier G. W. & K. Aziz. (1972). The flow of complex mixtures in pipes. Van Nostrand Reinhold Company, New York.

Graindor, M.J. (1964). Les tillites ante-cambriennes de Normandie. Geologische Rundschau, 54, 61–83.

Guan, B., Ruitang, W., Hambrey, M.B., Geng, W. (1986). Glacial sediments and erosional pavements near the Cambrian–Precambrian boundary in western Henan Province, China. Journal of the Geological Society, London, 143, 311–323.

Halverson, G.P., Maloof, A.C., & Hoffman, P.F. (2004). The Marinoan glaciation (Neoproterozoic) in Svalbard. Basin Research, 16, 297-324.

Halverson, G.P., Hoffman, P.F., Schrag, D.P., Maloof, A.C., & Rice, A.H.N. (2005). Toward a Neoproterozoic composite carbon-isotope record. Bulletin: Geological Society of America, 117, 1181–1207.

Hambrey M.J. & McKelvey, B.C. (2000). Major Neogene fluctuations of the East Antarctic ice sheet: Stratigraphic evidence from the Lambert Glacier region. Geology, 28(10), 887-890. Haughton, B.D.W., Barker, S., & McCaffrey, W. (2003). ‗Linked‘ debrites in sand-rich turbidite systems-origin and significance. Sedimentology, 50 459-482.

Haughton, P., Davis, C. & McCaffrey, W. (2006). Facies Prediction in Turbidite Fan Systems – Nature and Significance of ‗Linked Debrites' in Sand-Rich Versus Mixed Sand-Mud Systems" Recent Advances in Siliciclastic Facies Models: Implications for Reservoir Characterization II (SEPM), AAPG Annual Convention, Technical Program.

Hein, F.J. (1982). Depositional mechanisms of deep-sea coarse clastic sediments, Cap Enrag Formation, Quebec. Canadian Journal of Earth Science, 19, 267-287.

Heinrich, H. (1988). Origin and consequences of cyclic ice rafting in the Northeast Atlantic Ocean during the past 130,000 years, Quaternary Research, 29, 142-152.

147

Hiscott, RN. (1994). Traction-carpet stratification in turbidites- fact or fiction? Journal of Sedimentary Research, A64, 204-208.

Hoffman, P.F., Kaufman, A.J., Halverson, G.P., & Schrag, D.P. (1998). A Neoproterozoic Snowball. Earth Science, 281, 1342–1346.

Hoffman, P.F., & Schrag, D.P. (2000). Snowball Earth. Scientific American, 282, 62–75.

Hoffman, P. F., & Schrag D. P. (2002). The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14(3), 129-155.

Hoffmann, K.H., Condon, D.J., Bowring, S.A., & Crowley, J.L. (2004). U-Pb zircon date from the Neoproterozoic Ghaub Formation, Namibia: Constraints on Marinoan glaciation. Geology, 32, 817-820.

Hoffman, P.F. (2005). On Cryogenian (Neoproterozoic) ice-sheet dynamics and the limitations of the glacial sedimentary record. South African Journal of Geology, 108, 557–577.

Hoffman, P.F., Halverson, G.P., Husson, J.M., Hurtgren, M.T., & Schrag, D.P. (2006). Cap carbonate stratigraphy and ocean hyperstratification during meltdown of the Snowball/Slushball Earth in 635 Ma. Abstracts, Snowball Earth 2006, Ascona, Switzerland. p. 49.

Hoffman, P.F., & Li, Z.X. (2009). A palaeogeographic context for Neoproterozoic glaciation Palaeogeography, Palaeoclimatology, Palaeoecology, 277, 158–172.

Ineson, J.R., (1989). Coarse-grained submarine fan and slope apron deposits in a Cretaceous back-arc basin, Antarctica. Sedimentology, 36, 793– 819.

Johansson, M., & Stow, D.A.V. (1995). A classification scheme for shale clasts in deep water sandstones. In: A.J. Hartley & D.J. Prosser (Eds.), Characterization of Deep Marine Clastic Systems. Geological Society (London) Special Publication, 94, 221-241.

Kaye, C.A., & Zartman, R.E. (1980). A Late Proterozoic Z to Cambrian age for the stratified rocks of the Boston Basin, Massachusetts. In: D.R., Wones (Ed.), The Caledonides in the USA: Blacksburg, Virginia Polytechnic Institute and State University Memoir, 2, 257-261.

Kim, S.B., Chough, S.K., & Chun, S.S. (1995). Bouldery deposits in the lowermost part of the Cretaceous Kyokpori Formation, SW Korea: cohesionless debris flows and debris falls on a steep gradient delta slope. Sedimentary Geology, 98, 97–119.

Kirschvink, J.L. (2002). When All of the Oceans Were Frozen. Recherche, 355, 26–30.

148

Lahee, F.H. (1914). Late Paleozoic glaciation in the Boston Basin, Massachusetts. American Journal of Science, 187(4), 316-318.

LaForge, L. (1932). Geology of the Boston area, Massachusetts. U.S.Geological Survey, Bulletin, 839, p. 105.

Larsen, V., & Steel, R.J. (1978). The sedimentary history of a debris flow dominated alluvial fan -a study of textural inversion. Sedimentology, 25, 37-59.

Lenk, C., Strother, P.K., Kaye, C.A., & Barghoorn, E.S. (1982). Precambrian age of the Boston Basin: new evidence from microfossils. Science, 216, 619-620.

Lemon, N.M., Gostin, V.A. (1990). Glacigenic sediments of the late Proterozoic Elatina Formation and equivalents, Adelaide Geosyncline, South Australia. In: J.B., Jago, P.S., Moore (Eds.), The Evolution of a Late Precambrian–Early Palaeozoic Rift Complex: The Adelaide Geosyncline. Geological Society of Australia Special Publication, 16, pp. 149–163.

Lewis, J. P., Weaver, A. J., Johnston, S. T. & Eby M. (2003). Neoproterozoic ―snowball Earth‖: Dynamic sea ice over a quiescent ocean, Paleoceanography, 18 (4), 1092.

Lowe, D.R., & LoPiccolo, R.D. (1974). The characteristics and origins of dish and pillar structures. Journal of Sedimentary Petrology, 44, 484– 501.

Lowe, D.R. (1982). Sediment gravity flows: II. Depositional models with special reference to the deposits of high-density turbidity currents. Journal of Sedimentary Petrology, 52, 279-297.

Lowe, D.R. & Guy, M. (2000). Slurry-flow deposits in the Britannia Formation (Lower Cretaceous), North Sea: a new perspective on the turbidity current and debris flow problem. Sedimentology, 47, 31–70.

MacGabhann, B. A. (2005). Snowball Earth, the Dalradian Supergroup, and the Cryogenian- Ediacaran Boundary. 48th Irish Geological Research Meeting Programme and Abstracts, 33.

Martins-Neto, M.A. (1996). Lacustrine fan-deltaic sedimentation in a Proterozoic rift basin: the Sopa-Brumadinho Tectonosequence, southeastern Brazil. Sedimentary Geology, 106, 65-96.

Mawson, D., (1949). The Elatina glaciation. A third occurrence of glaciation

149 evidenced in the Adelaide System. Research Society of Southern Australia, 73, 117–121.

McMenamin, M.A.S., & Beuthin, J.D. (2008). Fine Clastics of the Boston Bay Group: New data and interpretations concerning depositional processes and environments. 21st Annual Keck Symposium, 209-212.

Middleton, G. V. (1967). Experiments on density and turbidity currents. III Deposition of sediments. Canadian Journal Earth Science, 4, 475-505.

Middleton, G.V., & Hampton, M.A. (1973). Sediment gravity flows: Mechanics of flow and deposition, In: G.V. Middleton, & A.H. Bouma, (Eds.), Turbidites and deep-water sedimentation: Los Angeles, Pacific section. Society of Economic Paleontologists and Mineralogists, 1-38.

Miller, N., Johnson, P.R., & Stern, R.J. (2007). Marine versus non-marine environments for the Jibalah Group, NW Arabian shield: A sedimentologic and geochemical survey and report of possible metazoan in the Dhaiqa Formation. The Arabian Journal for Science and Engineering, 33(1C), 56-77.

Moncrieff, A.C.M. and Hambrey, M.J. (1990). Marginal-marine glacial sedimentation in the late Precambrian succession of East Greenland. In: J.A. Dowdeswell, & J.D. Scourse (Eds.), Glacimarine Environments: Processes and Sediments. Geological Society London Special Publication, 387–410.

Mulder, T., & Alexander J. (2001). The physical character of subaqueous sedimentary density flows and their deposits. International Association of Sedimentologists, Sedimentology, 48, 269-299.

150

Murphy, J.B., van Staal, C.R., & Keppie, J.D. (1999). Middle to Late Paleozoic Acadian Orogeny in the northern Appalachians: A Laramide style plume-modified orogeny. Geology, 27, 653-656.

Nance, R.D., Murphy, J.B., Strachan, R.A., D‘Lemos, R.S., & Taylor, G.K. (1991). Late Proterozoic tectonostratigraphic evolution of the Avalonian and Cadomian terranes. Precambrian Research, 53, 41-78.

Nemec, W., Porebski, S.J., & Steel, R.J. (1980). Texture and structure of resedimented conglomerates: examples from Ksiaz Formation (Famennian-Tournaisian) south-western Poland. Sedimentology, 27, 519-538.

Nemec, W., Steel, R.J., Porebski, S.J., & Spinnangr, A. (1984). Domba Conglomerate, Devonian, Norway: process and lateral variability in a mass flow-dominated lacustrine fan- delta. In: E.H., Koster,& R.J. Steel (Eds.), Sedimentology of Gravels and Conglomerates. Canadian Society of Petroleum Geologists, 295–330, Calgary.

Nichols, G. (1999). Sedimentology & Stratigraphy. Blackwell Science, Osney Mead, Oxford. pp 62-75.

Nystuen, J. (1976). Late Precambrian Moelv tillite deposited on a discontinuity surface associated with a fossil ice wedge, Rendalen, southern Norway. Norsk Geologiske Tidsskrift, 56, 29–50.

Oard, M.J. (2008). An ancient ‗ice age‘ deposit attributed to subaqueous mass flow -again! Journal of Creation, 22(2), 36–39.

O‘Brien, S.J., Wardle, R.J., & King, A.F. (1983). The Avalon Zone: a Pan-African terrane in the Appalachian Orogen of Canada. Journal of Geology, 18, 195-222.

O‘Brien, S.J., O‘Brien, B.H., Dunning, G.R., & Tucker, R.D. (1996). Late Proterozoic and related peri-Gondwanan rocks of the Newfoundland Appalachians, In: R.D. Nance, & R.D. Thompson, (Eds.), Avalon Ian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado, Geological Society of America Special Paper, 304, 9-28.

Olcott, A.N., Sessions, A.L., Corsetti, F.A., Kaufman, A.J., & de Oliveria, T.F. (2005). Biomarker evidence for photosynthesis during Neoproterozoic Glaciation, Science, 310, 471- 474.

151

Passchier, S., & Erukanure, E. (2010). Pale environments and weathering regime of the Neoproterozoic Squantum ―Tillite‖, Boston Basin: no evidence of a snowball Earth. Sedimentology, 57(6), 1526-1544.

Peltier, W.R., Liu Y., & Crowley, J.W. (2007). Snowball Earth prevention by dissolved organic carbon remineralization, Nature, 450, 813–818.

Pe-Piper, G., & Murphy, J.B. (1989) Petrology of the Late Proterozoic Folly River Formation, Cobequid Highlands, Nova Scotia. Atlantic Geology, 25, 143-151.

Pollard, D. and Kasting, J.F. (2005) Snowball Earth: A thin ice solution with flowing sea glaciers. Journal of geophysical Research – Atmospheres, 110.

Postma, G., Nemec, W., & Kleinspehn, K.L. (1988). Large floating clasts in turbidites: a mechanism for their emplacement. Sedimentary Geology, 58, 47–61.

Potter, P.E., & Scheidegger, A.E. (1966). Bed thickness and grain size: graded beds. Sedimentology, 7, 233–240.

Poulsen, C.J. (2003). Absence of a runaway ice-albedo feedback in the Neoproterozoic. Geology, 31, 473-476.

Prave, A.R. (1999). Two diamictites, two cap carbonates, two δ13C excursions, two rifts: the Neoproterozoic Kingston Peak Formation, Death Valley, California. Geology, 27, 339–342.

Preiss,W.V. (2000). The Adelaide Geosyncline of South Australia and its significance in Neoproterozoic continental reconstruction. Precambrian Research, 100, 21–63.

Prior, D. B., Coleman, J. M., & Bornhold, B. D. (1982). Results of a known sea floor instability event. Geo-Marine Letters, 2, 117-122.

Pyles, D.R., & Jeanette, D.C. (2009). Geometry and architectural associations of co-genetic debrite-turbidite beds in basin-margin strata, Carboniferous Ross Sandstone (Ireland): Applications to reservoirs located on the margins of structurally confined submarine fans. Marine and Petroleum Geology, 26(10), 1974-1996.

Rankin, G. (1993). The volcanogenic Mount Rogers Formation and the overlying glaciogenic Konnarock Formation: Two late Proterozoic units in southwestern Virginia 26 p. U.S. G.P.O. (Washington, D.C. and Denver, CO)

152

Rast, N. & Skehan, J.W. (1983). The evolution of the Avalonian Plate. Tectonophysics, 100, 257-286.

Rehmer, J.A. & Roy, D.C. (1976). The Boston Bay Group: the boulder bed problem, In: B. Cameron (Eds.), Geology of southeastern New England: Princeton, New Jersey, New England Intercollegiate Geologic Conference, 68th Annual Meeting, Guidebook, 71-91.

Rocheleau, M. & LaJoie, J. (1974). Sedimentary structures in resedimented conglomerate of the Cambrian flysch, L‘Islet, Quebec Appalachians. Journal of Sedimentology and Petrology, 45, 3-45.

Rossetti, D.F., (1999). Soft-sediment deformation structures in late Albian to Cenomanian deposits, Sao Luis Basin, northern Brazil: evidence for palaeoseismicity. Sedimentology, 46, 1065–1081.

Runnegar, B. (2000). Loophole for the Snowball Earth. Nature, 405, 403-404.

Sadler, P. M. (1982). Bed-thickness and grain size in turbidities. Sedimentology, 29, 37-51.

Sayles, R.W., & LaForge, L. (1910). The glacial origin of the Roxbury Conglomerate. Science, 32, 723-724. Sayles, R.W. (1914). The Squantum Tillite. Bulletin Harvard Museum Comparative Zoology, 66, 141-175.

Sayles, R.W. (1919). Seasonal deposition in the aqueo-glacial sediments. Memoirs of the Museum of Comparative Zoology, Harvard College, 47, 67 p.

Schermerhorn, L. J. G. (1974). Late Precambrian mixtites: Glacial and/or non-glacial? American Journal of Science, 274, 673–824.

Schrag, D.P., Berner, R.A., Hoffman, P.F. & Halverson, G.P. (2002). On the initiation of a snowball Earth. Geochemical Geophysics Geosystems, 3, 1029-1036.

Schuppers, J.D. & Martinius, A.W. (1994). Modified-grain-flow deposits from the upper Cretaceous Vallcarga Formation, south-central Pyrenees, Spain. Journal of Sedimentary Research A64, 518–523. Schwab, F.L. (1981). Late Precambrian tillites of the Appalachians, In: M.J. Hambrey, & W.B. Harland, (Eds.), Earth‘s Pre-Pleistocene Glacial Record. Cambridge University Press, 751-755.

Schwarz, H.U. (1982). Subaqueous slope failures-experiments and modern occurrences. Contributions to Sedimentology, 11, 116.

Shanmugam, G. (2000). 50 years of the turbidite paradigm: deep-water processes and facies models- a critical perspective. Marine and Petroleum Geology, 17, 285-342.

Shanmugam, G. (2002). Ten turbidite myths. Earth-Science Reviews, 58, 311–341.

153

Shen, B., Xiao, S., Dong, L., Zhou, C. & Liu, J. (2007). Problematic macrofossils from Ediacaran successions in the North China and Chaidam blocks: implications for their evolutionary roots and biostratigraphic significance. Journal of Paleontology, 81, 1396-1411.

Smith, G.W. & Socci, A.D. (1990). Late Precambrian sedimentary geology of the Boston Basin, In: A. D. Socci, S.J. Skehan, & G.W Smith (Eds.), Geology of the composite Avalon terrane of southern New England. Geological Society of America Special paper, 245, 75-84.

Socci, A.D. & Smith, G.W. (1987). Recent sedimentological interpretations in the Avalon terrane of the Boston Basin, Massachusetts: Maritime Sediments and Atlantic. Geology, 23, 13-39.

Socci, A.D. & Smith, G.W. (1990). Stratigraphic implications of facies within the Boston Basin. In: A.D., Socci, J.W. Skehan, & G.W. Smith, (Eds.), Geology of the Composite Avalon terrane of southern New England: Boulder, Colorado, Geological Society of America Special Paper, 245, 55-74.

Sohn, Y.K., Choe, M.Y., & Jo, H.R. (2002). Transition from debris flow to hyperconcentrated flow in a submarine channel (the Cretaceous Cerro Toro Formation, southern Chile). Terra Nova 14, 405–415.

Steel, R.J. & Thompson, D.B. (1983). Structures and textures in Triassic braided stream conglomerates (‗Bunter‘ Pebble Beds) in the sherwood Sandstone Group, North Staffordshire), England. Sedimentolgy, 30, 341-368.

Stern, R.J., Avigad, D., Miller, N.R. & Beyth, M. (2006). Geological Society of Africa Presidential Review: Evidence for the Snowball Earth Hypothesis in the Arabian-Nubian Shield and the East African Orogen. Journal of African Earth Sciences, 44, 1–20.

Sylvester, Z. & Lowe, D.R. (2004). Textural trends in turbidites and slurry beds from the Oligocene flysch of the East Carpathians, Romania. Sedimentology, 51, 945–972.

Talling, P.J., Amy, L.A., Wynn, R.B., Peakall, J. & Robinson, M. (2004). Beds comprising debrite sandwiched within co-genetic turbidite: origin and widespread occurrence in distal depositional environments. Sedimentology, 51, 163– 194.

Thomas, G.S.P. & Connell, R.J. (1985). Iceberg drop, dump, and grounding structures from Pleistocene glacio-lacustrine sediments, Scotland. Journal of Sedimentary Petrology, 55, 243- 249.

Thompson, M.D. (1993). Late Proterozoic stratigraphy and structure in the Avalonian magmatic arc southwest of Boston Massachusetts. American Journal of Science, 293, 725-753.

Thompson, M.D., Hermes, O.D., Bowring, S.A., Isachsen, C.E., Besancon, J.R. & Kelly, K.L. (1996). Tectonostratigraphic implications of Late Proterozoic U-Pb zircon ages in the Avalon Zone of southeastern New England. In: R.D. Nance & M.D. Thompson (Eds.), Avalonian and

154

Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America Special Papers, 304, 179-191.

Thompson, M.D. & Bowring, S.A. (2000). Age of the Squantum ―tillite‖, Boston Basin, Massachusetts: U-Pb zircon constraints on terminal Neoproterozoic glaciation. American Journal of Science, 300, 630-655.

Thompson, M.D. & Grunow, A.M. (2004). A Roxbury Review: conglomerates of greater Boston, Massachusetts, In: L.S. Hanson, (Ed.), Guidebook to field trips in Rhode Island and Massachusetts: Salem, MA., 96th Annual Meeting of the New England Intercollegiate Field Conference, p. C2-1 - C2-21.

Thompson, M.D., Grunow, A.G. & Ramezani, J. (2007). Late Neoproterozoic paleogeography of the Southeastern New England Avalon Zone: Insights from U-Pb geochronology and paleomagnetism. Bulletin: Geological Society of America, 119, 681-696. Trench, A. & Torsvik, T.H. (1992). The closure of the Iapetus Ocean and Torniquist Sea: New Paleomagnetic Constraints. Geological Society of London, 149, 867-870.

Trop, J.M., Ridgway, K.D., Sweet, A.R. & Layer, P.W. (1999). Submarine fan deposystems and tectonics of a Late Cretaceous forearc basin along an accretionary convergent plate boundary, MacColl Ridge Formation, Wrangell Mountains, Alaska. Canadian Journal of Earth Science, 36, 433–458.

Walker, R.G. (1975). Generalized facies model for resedimented conglomerates of turbidite association. Bulletin: Geological Society of America, 86, 737–748.

Walker, R.G. (1984). Shelf and Shallow Marine Sands. In: Facies Models (E.) R.G. Walker, 2nd edn, pp. 141–170. Geological Association of Canada, Toronto.

Williams, H. & King, A. F. (1979). Trepassey map area, Newfoundland. Canadian Geological Survey Memoirs, 389, 24 pp.

Williams, G. & Schmidt, P. (2000). Proterozoic equatorial glaciation: Has 'snowball Earth' a snowball's chance? The Australian Geologist, 117, p. 21.

Williams, J.A. (2008). Laminates and dropstones in the Cambridge Argillite (Ediacaran), Hewitt‘s Cove, Hingham, Massachusetts, 21st Keck Geology Consortium, 234-237. Winterer, E. (1963). Late Precambrian pebbly mudstone in Normandy, France. In: A.E.M. Nairn (Ed), Problems in Palaeoclimatology, 159-178, Interscience, London.

Wolfe, W. (1976). Geology of Squaw Head, Squantum, Massachusetts In: B. Cameron, (Eds.) Geology of southeastern New England: Princeton, New Jersey, New England Intercollegiate Geologic Conference, 68th Annual Meeting, Guidebook, 107-116.

155

Xiao, S., Bao, H., Wang, H., Kaufman, A.J., Zhou, C., Li, G., Yuan, X. & Ling, H. (2004). The Neoproterozoic Quruqtagh Group in eastern Chinese Tianshan: evidence for a post-Marinoan glaciation. Precambrian Research, 130, 1–26.

Zaleha, M.J. & Wiesemann, S.A. (2005). Hyperconcentrated flows and gastroliths: sedimentology of diamictites and wackes of the Upper Cloverly Formation, Lower Cretaceous, Wyoming, U.S.A. Journal of Sedimentary Research, 75(1), 43–54.

Zartman, R.E. & Naylor, R.S. (1984). Structural implications of some radiometric ages of igneous rocks in southeastern Massachusetts. Bulletin: Geological Society of America, 95, 522- 539.

Zhao, X., Zhang, L., Zou, X., Wang, S. & Hu, Y. (1980). Sinian tillites in Northwest China and their stratigraphic significance. In: Tianjin Institute of Geology and Mineral Resources (Ed.), Research on Precambrian Geology Sinian Suberathem in China. Tianjin Science and Technology Press, Tianjin, pp. 164–185.

Zhao Y.Y., Zheng Y.F. & Chen, F.K. (2009) Trace element and strontium isotope constraints on sedimentary environment of Ediacaran carbonates in southern Anhui, South China. Chemical Geology, 265, 345-362.

Zhao Y.Y. & Zheng Y.F. (2010). Stable isotope evidence for involvement of deglacial meltwater in Ediacaran carbonates in South China. Chemical Geology, 271, 86-100.

156

CHAPTER 4:

MASS FLOW REWORKING OF GLACIAL SEDIMENT ON MOUNT RAINIER: IMPLICATIONS FOR UNDERSTANDING THE FATE OF TERRESTRIAL GLACIAL SEDIMENT ON NEOPROTEROZOIC ISLAND ARCS

ABSTRACT

During the short-lived Neoproterozoic Gaskiers glaciation (c. 580 Ma) glaciers formed on the summits of volcanoes scattered along island arcs of the Avalonian-Cadomian Terranes then located around the margin of Gondwana (peri-Gondwanan terranes). At this time these terranes were composed of several volcanoes and marginal arc-related rift basins. Glaciation along the terranes is recorded in a single rift basin by rare ‗ice rafted‘ striated and faceted clasts in the deep marine strata of the Gaskiers Formation in Newfoundland, Canada (Avalon

Terrane). As a result there is a growing consensus that glaciers were local and topographically restricted to high altitude volcanoes along the Avalonian-Cadomina Terranes. The presence of glacial sediment in the Gaskiers Formation is interpreted to record the growth of these glaciers to sea level. On-land terrestrial glacial facies (e.g., tillites) have not been identified anywhere along the Avalon-Cadomian Terranes. To better understand the fate of glacial sediments and landforms on the landward margins of ancient glaciated volcanic arc basins this study examines post-depositional processes reworking glacial sediment on Mount Rainier, the highest glaciated stratovolcano in the Cascade Range of Washington, USA. Mount Rainier supports some 100 km2 of glacier ice and snow including several large valley glaciers, and as such represents an excellent analogue for terrestrial glacial settings along the Avalonian-

Cadomian Terranes. The results of this study show that primary glacial sedimentary deposits

157 do not survive on the steep volcanically active and earthquake-prone slopes of the volcano that experience the abrupt melt of glacier ice during eruptions, heavy rainfall and spring snowmelt runoff. Diamict deposits on Mount Rainier that have been widely mapped as ‗till‘ or ‗glacial drift‘ are shown here to be debris flow deposits. One exceptionally voluminous (4 km3; 200 km2) volcanogenic lahar (‗Osceola Mudflow‘; 5.6 ka) described earlier as a ‗till sheet‘ has been reinterpreted as a massive mudflow deposit recording the collapse of the upper glaciated flank of Mount Rainier. On Mount Rainier, primary glacial sediment fingerprinted by subglacially- shaped and striated clasts, is continuously reworked and ‗diluted‘ by large volumes of bouldery rock fall debris, pyroclastic debris and ash during downslope flow. Large outburst floods triggered by rapid melting of ice during eruptions also rework glacial debris as

‗hyperconcentrated flows‘ destroying any primary glacial signal. In the long term, the sedimentary record of glaciation on Mount Rainier is essentially ‗invisible‘ as a consequence of downslope reworking. Andesitic lavas with rapidly-chilled ‗ice contact‘ margins provide rare records of glacier ice on the volcano.

The value of these observations is that they identify the depositional processes that challenge the ability of geologists to recognize terrestrial glacial deposits of ancient glaciated and volcanically influenced settings such as the Gaskiers Formation.

158

4.1 INTRODUCTION AND PURPOSE OF STUDY

True glacial tillites and other ice-contact facies deposited below the margins of ice sheets and glaciers are rarely preserved in the pre-Pleistocene glacial record, which is biased toward selectively preserved marine glacial diamictites (Eyles, 1993; Crowell, 1999). In the

Neoproterozoic, many diamictites in glacially influenced marine basins record downslope mass flow and the reworking of primary glacial sediments as debrites; the same debrite facies, however, are produced in non-glacial environments by the mixing of gravel and mud on unstable slopes in tectonically active basins. In contrast to the Snowball Earth Hypothesis

(SEH; Hoffman et al., 1998; Evans, 2000; Hoffman & Schrag, 2002; Kirschvink, 2002; Schrag et al., 2002; Rice et al., 2003; Hoffman, 2005), facies analysis studies support an emerging alternative model that emphasizes the importance of active tectonics during the formation of rift basins as Rodinia broke apart (c. 750-600 Ma) and the production of large volumes of non- glacial diamictites by mass flow processes notably slumping and debris flow within incipient rift basins or along the flanks of volcanic arcs (Schermerhorn, 1974, 1983; Eyles and Eyles,

1983; Eyles, 1993; Young, 1995, 2002; Crowell, 1999; Eyles and Januszczak, 2004; Eyles et al., 2007).

The Gaskiers glaciation at ~580 Ma (Bowring et al., 2003) is based on the type site in

Newfoundland, Canada, and is characterized by the clustering of glaciers on the peaks of volcanoes along the outer rim of Gondwana (Avalonian-Cadomian Orogenic Belt). Direct terrestrial sedimentological evidence for glaciation, such as those observed around modern ice margins (e.g., tillites) has not been found anywhere along this Belt, yet evidence of a glacial influence on sedimentation is present in a single case, the Gaskiers Formation of

Newfoundland, in the form of striated and faceted dropstones within the marine facies of the

159 basin (Bruckner and Anderson, 1971; Williams and King, 1979; Eyles and Eyles, 1989; Eyles and Januszczak, 2004; Eyles, 2008; Carto and Eyles, 2011).

To better understand the absence of a terrestrial glacial record along the Avalon

Terrane, in regards to the Gaskiers Formation, a detailed field study was conducted on the sedimentary environments and deposits of Mount Rainier, a large modern glaciated stratovolcano in Washington State, USA. The objective of field study was to gain insight into the preservation potential of glacial sediment on the landward margins of ancient glaciated volcanic arc basins. Mount Rainier was chosen for study as it presents an excellent laboratory for studying the interaction of glaciation and volcanism. Over the last 10 ka alone several lahars and glacial outburst floods have been generated from the volcano and as a result, have reworked glacial sediment of varying ages (between 130-6 ka) down the valleys of Mount

Rainier. The depositional products of these flows presently cover the flanks of the volcano and its lowland areas (Puget Sound Lowland) (Fiske et al., 1963; Crandell, 1971; Scott et al., 1995;

Pringle, 2008).

The main objectives of this study are: (1) examine and describe lithofacies of the various diamict successions generated at Mount Rainier, (2) to identify the dominant depositional processes for the diamict deposits, and (3) investigate the change in deposit-type with distance away from the glacial source in order to discern the influence of reworking on the preservation of the sedimentary properties of primary glacial till (e.g., subglacially shaped and striated clasts).

160

4.2 PHYSICAL SETTING AND SEDIMENTARY DEPOSITS OF MOUNT RAINIER

Active glaciated volcanoes are very common around the Pacific Rim in southern

Alaska, the Aleutian Chain and in southern Chile. Mount Rainier in Washington State, USA, was chosen for this study because Mount Rainier‘s deposits are readily accessible and have been extensively mapped. Mount Rainier is a relatively young (< 1 Ma; Wood and Kienle,

1990; Sisson, 1995) andesitic stratovolcano in the Cascade Range, which results from the subduction of the Juan de Fuca, Explorer, and the Gorda Plates (remnants of the much larger

Farallon Plate) under the western margin of the North American Plate (Crandell, 1963;

Crandell and Miller, 1974; Sisson and Lanphere, 1997) (Fig. 4.1). Mount Rainier is the highest

(4392 m) volcano in the Cascade Range and experiences frequent explosive phreatomagmatic eruptions due to magma-ice interactions, sector collapse and large debris flows of volcanic and glacial sediment (e.g., McBirney, 1968; Green et al., 1988; Pringle, 2008). As a result of its height and proximity to Pacific moisture, Mount Rainier supports several large valley glaciers

(e.g., Nisqually, Carbon and Emmons glaciers), smaller cirque glaciers (e.g., Paradise glacier) and snowfields, which collectively cover 92 km2 of the volcano with a total ice volume of about 4.4 km3 (Driedger and Kennard, 1986; Walder and Driedger, 1993). Primary glacial sediment on Mount Rainier is represented by ‗tills‘ and ‗drift‘ deposits left by several

Pleistocene glacial advances, such the Wingate Hill Drift (600-300 ka), Hayden Creek (35–50 ka), Evans Creek Drift (25-15 ka) and McNeeley Drift (11 ka) (Fig. 4.2) (Crandell, 1969, 1971;

Crandell and Miller, 1974). ‗Neoglacial‘ re-advances occurred on the volcano during the

Holocene at 6.6 ka, 2.8 ka and 2.6 ka (Crandell and Miller, 1974; Pringle, 2008). During the

‗Little Ice Age‘ event at about 1750-1890 A.D, glaciers on Mount Rainier expanded once again

161

(the Garda Drift 0.7 ka; Crandell and Miller, 1974) and have since retreated about 3 km up valley. Ongoing deglaciation has left prominent lateral and terminal moraines and glaciofluvial outwash along valley floors. Glacial sediments are dominated by supraglacial sediment derived from rockfalls above the snowline with much lesser amounts of subglacially derived sediment

(e.g., Mills, 1979).

The present summit dome of Mount Rainier consists of andesitic to dacitic lava flows, and welded and non-welded block- and ash-flow tuffs (Crandell and Miller, 1974), and is built on pre-Rainier granodiorite of the Miocene Tatoosh pluton and older middle Tertiary volcanic rocks (Crandell, 1963; Fiske et al., 1963; Vance et al., 1987; Vallance and Scott, 1997). Highly explosive eruptive events occurred during the last (Wisconsin) glaciation, clustering around l30 to 90 ka and between 70 to 30 ka (Wood and Kienle, 1990). Eruptions from Mount Rainier continued into the Holocene marked by eleven easily recognizable pumiceous tephra layers

(vesicle-rich and vesicle-poor), eight of which fall between 6.5 and 4.0 ka, at 1.0 and 2.3 ka and one between 1820 and 1854 A.D (Mullineaux, 1974).

Evidence of ice-magma interaction is found in the widespread presence of columnar jointed andesite and dacite lava flows that form steep-sided ridges separating the major valleys that radiate from the volcano‘s summit (Fig. 4.3) (Lescinsky and Sisson, 1998). These rocks were originally interpreted as remnants of extensive flows that were later cut by deep glacial erosion to produce valleys (Fiske et al. 1963). However, the glassy textures and young ages (<

40 ka) of these rocks indicate an ice-contact origin (Lescinsky and Sisson, 1998; Pringle, 2008) when lava flowed along the margin of valleys that were inundated by glaciers. Following deglaciation, lava deposits were left perched on ridge tops (Mathews, 1952). This process is also reported from Mount Garibaldi in British Columbia (Mathews, 1952), Mount Hekla in

162

Iceland, and Villarrica and Hudson volcanoes in Chile (Major and Newhall, 1989; Naranjo et al., 1993).

Holocene (< 10 ka) sedimentary processes at Mount Rainier are dominated by extensive lahars (volcanic mudflows) (Crandell, 1963; Crandell and Miller, 1974; Scott et al.,

1992). More than 60 large lahars have been generated at Mount Rainier extending more than

100 km from the volcano into Puget Sound Lowland. Major mudflows include the Osceola (5.6 ka), Round Pass (2.6 ka) and Electron (~A.D. 1502–1503) mudflows and the Paradise Lahar

(5.8-6.6 ka) (Fig. 4.4) (McBirney, 1968; Crandell 1969; Swanson et al. 1989; Wood and

Kienle, 1990; Scott et al. 2001). The Osceola Mudflow is the most extensive covering 350 km2 filling the White River valley to thicknesses of up to 150 m (Crandell, 1971; Vallance and

Scott, 1997). The Round Pass Mudflow in the Puyallup River valley is characterized by its great thickness (locally up to 250 metres; Crandell, 1971). The National Lahar is a relatively small event, which inundated the Nisqually River valley to thicknesses of up to 40 metres. The

Electron Mudflow reached the Puget Lowland extending about 60 km down valley (Crandell,

1963). Both the Paradise Lahar and the Osceola Mudflow were triggered by eruptions whereas the Round Pass and Electron Mudflows and National Lahars are not linked to any known eruptions (Swanson et al. 1989; Wood and Kienle, 1990).

Volcanism coeval with episodes of expanding Neoglacial ice cover has also produced numerous glacial outburst floods (―jökulhlaups‖; Walder and Driedger, 1993) (Fig. 4.2); other flood events are triggered by heavy rainfall and rapid spring snowmelt (Frank, 1985). Flood waters moving down valley evolve into large debris flows by incorporating large volumes of glacial and volcaniclastic sediment (Driedger and Fountain, 1989; Scott et al., 1992).

163

4.3 STUDY AREA, METHODS AND TERMINOLOGY

It became immediately apparent during the initial stages of fieldwork that existing published terminology used to classify surficial sediment on Mount Rainier is simplistic.

Terms such as ‗till‘ and ‗bouldery till‘ (Crandell, 1971; Crandell and Miller, 1974) have been liberally used for any poorly sorted deposit composed of a mixture of glacial sediment, bedrock rubble, pyroclastic debris and coarse-grained proximal fluvial deposits which has been reworked and mixed together by mass flow (e.g., Boulton and Eyles, 1979). Other workers

(Driedger, 1986; Walder and Driedger, 1993) have similarly recognized the heterogeneous nature of deposits and employ the umbrella term ‗glacial drift‘ in recognition of widespread mixing of ‗till-like‘ diamict, rubbly debris flow facies and ‗outwash.‘ Even here however, the term ‗glacial‘ disguises the importance of mass flow and the minimal influence of ice.

This study presents facies descriptions of deposits previously mapped as 1) ‗till‘ and

‗glacial drift‘ deposits left by both valley glaciers and continental ice that invaded the Puget

Lowland around Mount Rainier; 2) lahar deposits (Osceola Mudflow) and; 3) outburst flood deposits (Figs. 4.2, 4.4). The Evans Creek (~25-15 ka) and Hayden Creek (~35-50 ka) till/drift deposits record the readvance of alpine glaciers during the last (Wisconsin) glaciation (known locally as the Fraser glaciation) when valley glaciers extended as far as 25 km beyond present glacier termini (Crandell and Miller, 1974). In contrast, Vashon Till was deposited by the

Puget lobe of the Cordilleran ice sheet as it expanded south into the Fraser and Puget Sound

Lowlands between 22 ka and 15 ka (Waitt and Thorson, 1983). The geographic scope of the study was narrowed by focusing on outcrops exposed within and around the valley of the

White River (Figs. 4.2,4.4) where till/drift deposits are well exposed (Driedger, 1986), as well as deposits of the extensive and thick Osceola Mudflow. Exposures of glacial outburst flood

164 deposits were examined in the Kautz, Tahoma and Van Trump Park Creeks, which drain the

Kautz, South Tahoma, and Van Trump alpine Glaciers, respectively (Fig. 4.2).

Sedimentological and stratigraphic data were collected and lithofacies documented using standard facies logging procedures (Miall, 1978; Eyles et al., 1989; Nicholas, 1999). This work involved making detailed observations of facies types and associations, lateral, longitudinal, and vertical gradations in the deposits, bedding features (contacts, thickness, and geometry), lithology types, particle-size distribution, and clast features (shape10, size11, lithology, and texture). Emphasis was placed on documenting the down valley trend in the abundance of striated and faceted clasts, as well as clast shape found in till/drift, lahar, and outburst flood deposits.

4.4 STUDY RESULTS

4.4.1 Glacial till/drift

Two types of glacial sediment were examined; that from local valley glaciers and that from continental ice that invaded the Puget Lowlands surrounding Mount Rainier.

Thick (up to 10 m) poorly sorted diamict deposits linked to local valley glaciers (Evans Creek

Drift) were examined in the Nisqually River valley (locality C; Fig. 4.2) and the Van Trump

River valley (locality D; Fig. 4.2), which occur near the terminus of the Nisqually Glacier and

Van Trump valley glaciers. The deposits are a grey-purplish coloured, matrix-supported massive diamict composed of a mixture of randomly orientated, angular to subangular clasts of andesite that range from pebble- to-boulder-sizes supported by a coarse silty sand matrix. The

10 Clast roundness was determined using the visual roundness chart of Krumbein (1941). 11 Clast size is based on measurements of the c-axis length of the clasts (the longest axes of the clast). 165 diamict at locality C is overlain by a thin bed (few centimeters thick) of the Paradise Lahar (a yellowish-orange bed of angular to subrounded rock fragments in a matrix consisting of sand, silt, and clay) and pumice-rich tephra layers (< 1cm thick) derived from Mount Rainier (Figs.

4.5A,B and 4.9A).

At a more intermediate position along the White River valley (locality E; Fig. 4.2), a 9- m-thick, slightly weathered diamict composed of subangular andesitic clasts ranging from gravel-to boulder-sized is exposed. This deposit is massive and has a coarse sandy-silt matrix

(Fig. 4.5C). This exposure has been interpreted as ‗ablation till‘ of Evans Creek age (Pringle,

2008). At a more distal exposure at (locality F; Fig. 4.2), a thick bed of Evans Creek Drift (5 m thick), lithologically similar to the bed exposed at the two previous localities, overlies a glacially polished rhyolitic sill. Clasts are noticeably more rounded (ranging from subrounded- subangular) and range in size from pebbles to boulders and are dominantly andesitic in composition (Figs. 4.5D,E and 4.9B). At locality G (Fig. 4.2) another outcrop of Evans Creek

Drift is exposed consisting of a thick (12 m) diamict resting on basalt. This deposit has been mapped as terminal morainal material representing the farthest advance of valley glaciers

(Campbell, 1975).

Both matrix and clast lithologies, as well as clast shape are similar to locality F, but better sorting, subrounded clasts, and well-defined normal grading characterize this deposit. Clast size ranges from pebble-to cobble-sizes. Clasts are randomly orientated and unlike the previous deposits, this bed is characterized by a lack of boulders (Figs. 4.5F, G and 4.9C).

The most distal deposit of Evans Creek Drift occurs at locality H (Fig. 4.2). This diamict is 4 metres thick, massive and matrix-supported, and is lithologically similar to the other deposits of the Evans Creek age. Glacial faceting of clasts is common. Clasts are pebble-

166 to cobble-sized and are subrounded in shape composed of multi-lithic clasts including andesite, granodiorite and pyroclastic debris (Figs. 4.5H-J).

No striated clasts were identified in any of the Evans Creek glacial drift deposits, but many angular clasts were identified in upper valley deposits typical of debris derived from rock falls and carried supraglacially on the ice surface. A statistical analysis of the relationship between the degree of clast roundness and changes in elevation was conducted on the Evans

Creek glacial drift deposits. Figure 4.6 shows a scatter-plot of the data with a ‗best-fit‘ linear regression line. The results show that there is a strong (and causal) relationship between these two variables (r= 0.91) indicating that there is a strong (and predictable) down valley trend towards increased rounding of clasts in drift deposits of Evans Creek.

Hayden Creek Drift at locality I (Fig. 4.2) occurs as a 7-m-thick matrix-supported diamict composed of pebble- to boulder-sized clasts. Clasts are supported in a yellowish-brown coarse sand matrix; they are typically subrounded to subangular and composed of andesite and granodiorite (Figs. 4.7A, B). A distal deposit of the Hayden Creek Drift (locality J; Fig. 4.2) is a matrix-supported diamict (3.5 m thick) composed of a mixture of cobble-to-boulder sized clasts that are predominantly subrounded in shape; the proportion of coarse gravel-to-pebble sized clasts (< 15 cm) are subangular in shape. Clasts are supported by a fine-to medium- grained sand matrix. Significantly, no striated or facetted clasts were identified.

Outcrops of Vashon Till examined in the Ohop Valley (localities A and B; Fig. 4.2) are massive and crudely stratified matrix-supported diamicts up to 12 m thick, with subangular clasts up to boulder sizes. At this locality, deposits of Vashon Till overlie lahar and lahar-run- out deposits. Clast lithologies include a mixture of andesite, granodiorite, basalt, and diorite,

167 supported in a light grey-coloured coarse sand matrix, many of which are striated. (Figs. 4.8A,

B and Fig. 4.9D).

4.4.2 Osceola Mudflow

The 5.6 ka-old (4800 radiocarbon yr; Crandell, 1971) Osceola Mudflow is an exceptionally large lahar deposit by any global standard. It transported at least 3.8 km3 of rock debris covering more than 200 km2 of the Puget lowland (Crandell and Waldron, 1956;

Vallance and Scott, 1997). The mudflow developed when the northeast side of Mount Rainier‘s edifice collapsed and avalanched down the White River valley out into the Puget Sound

Lowland area 5.6 ka (Crandell, 1971). The source area of the mudflow is now hidden beneath the broad upper part of the Emmons and Winthrop Glaciers and possibly includes as much as

600 m of volcanic rock from the former summit region of Mount Rainier (Crandell, 1963,

1971). The avalanche transformed downslope within 2 km of its source into a matrix-rich debris flow by incorporating water, pyroclastic debris and ash, alluvium, glacial outwash and till (Fiske et al., 1963; Vallance and Scott, 1997).

The lahar is more than 100 km in length and comprises a conspicuous distinctly orange- coloured diamict along the White River valley (Fig. 4.4). The lahar had a velocity of ~19 m/s and maximum inferred instantaneous discharge of some 2.5 x 106 m3/s (Vallance and Scott,

1997). Crandell and Waldron (1956) identified charred wood, and the presence of crudely developed normal-grading and possible cavities left by the melt of ice entrained within the flow. Outcrops of the mudflow typically form 5-20 m thick fills along valley bottoms and < 3m thick covers on valley sides as much as 200 m above the present river level (Vallance and

168

Scott, 1997). The most extensive outcrops occur in Glacier Basin and in the Inter Fork valley; the most proximal outcrops (upper limit of inundation) of the mudflow occur on the ridge tops of Glacier Basin (400 m above valley bottom) and on the valley floor of Glacier Basin (Pringle,

2008). In Glacier Basin (locality K; Fig. 4.4), the Osceola Mudflow occurs as a massive, matrix-supported diamict composed of angular to subangular cobble-to- large boulder-sized clasts (< 2 m in diameter) of andesite and altered andesite. Beds range in thickness from 1 to 20 m. Clasts lack a consistent orientation and are supported by a distinctly orange-coloured sand- silt-clay matrix, derived from deeply weathered clayey volcanic ash and debris (Crandell and

Waldron, 1956). Striated and polished boulders are common (35 % of clasts examined) (Figs.

4.10A-D and 4.13A).

In the Inter Fork valley (9 km down valley of Glacier Basin at locality L, Fig. 4.4), a thick (150 m) deposit of the Osceola Mudflow occurs along the south bank of the valley wall ~

400 m above the valley bottom (Fig. 4.11A). The mudflow at this locality is massive, matrix- supported and composed of subangular cobble-to- large boulder-sized clasts (up to 3 m in diameter) of andesite and altered andesite. The inspection of clasts for glacial striae or shaping was prevented due to the confinement of this deposit to the high valley walls. Some 14 km down valley from Glacier Basin, a 24-m-thick bed of the Osceola Mudflow, lithologically and texturally similar to its upslope equivalents, is well exposed along the Sunrise Road (locality

M; Figs. 4.4, 4.11B). Striated cobbles were identified in deposits at this locality (15 % of clasts examined). In contrast, ~22 km down valley from Glacier Basin, the mudflow is significantly different as a result of flow transformation (locality N; Fig. 4.4), comprising a much more clay- rich matrix and higher frequency of rounded clasts, which range in size from pebbles to cobbles. Clasts are also multi-lithic consisting of andesite, altered andesite granodiorite and

169 pumice. Striated pebbles were identified in deposits at this locality (15 % of clasts examined)

(Figs. 4.11C, D and Fig. 4.13B).

Distal exposures of the Osceola at locality O (Fig. 4.4) occur near the town of

Greenwater ~56 km down valley from Glacier Basin. Here, the mudflow consists of massive and reversely graded beds (2-8 m thick) of matrix-supported diamict composed of subrounded clasts of andesite supported in a clay-rich matrix measuring up to 17 m thick, interbedded with thick (1-4 m) beds of massive sand and gravel. Striated clasts are still present (5 % of clasts examined) (Figs. 4.11E-H and 4.13C). The most distal deposit of the mudflow (locality P; Fig.

4.4) at Mud Mountain Dam (~70 km from Glacier Basin) shows a 24 m-thick diamict succession resting unconformably on Vashon glacial outwash gravels (Pringle, 2008). Diamict displays crudely developed bedding generated by superimposition of matrix-supported, massive diamict composed of cobble-to boulder-sized clasts that are subrounded to rounded in shape (Fig. 4.11J). The clasts within this deposit lacked glacial striae.

The degree of clast roundness and the abundance of striated clasts in the Osceola

Mudflow deposits were each measured and plotted as a function of elevation. The scatter-plots of the data with their corresponding linear regression lines are presented in Figures 4.12A/B.

The data shows that it can be assumed that there is a causal relationship between these two variables and a strong (and predictable) down valley trend towards increased clast roundness

(r= 0.94; Fig. 4.12A) and a reduction in the abundance of striated clasts (r= 0.91; Fig. 4.12B) in deposits of Osceola Mudflow.

170

4.4.3 Outburst flood deposits

Repeated eruptions of Mount Rainier are associated with the catastrophic melt of summit ice covers and the repeated reworking of glacial sediments downslope as outburst floods (‗jokulhlaups‘; Walder and Driedger, 1993). Glacial outburst flood deposits are well exposed in the Kautz Creek, Tahoma Creek and Van Trump Creek valleys, which drain the

Kautz, South Tahoma, and Van Trump Glaciers respectively (Fig. 4.2). The Kautz Creek flow event is the largest and youngest (circa 1947). Eyewitness reports indicate that the flood began when the lower 1.6 km of the Kautz Glacier collapsed (Richardson, 1968; Crandell, 1971) gouging a 91-m-deep ice-walled canyon and eroding a 18-m-deep swath of glacial gravels, as it moved down the Kautz Creek valley, forming a series of bouldery debris flows (Crandell and

Mullineaux, 1967; Richardson, 1968; Crandell, 1971). Some 40 million m3 of sediment was moved during flow, including boulders up to 4 m in diameter. Erosion by Kautz Creek into these deposits has exposed the deposits of the 1947 event along the valley.

At least 23 outburst flood flows have been generated in Tahoma Creek (Fig. 4.2) since

1967 carving a deep gorge into sediments and stagnant ice at the terminus of South Tahoma

Glacier. Most recently, in 2001, outburst flood flows moved down Van Trump Creek due to the sudden release of meltwater from the glacier (Vallance et al., 2003). Floodwaters ‗bulked up‘ as they crossed and cut into large areas of stagnant, debris-rich ice (Pringle, 2008) generating a series of debris flows that spilled into the Nisqually River (Vallance et al., 2003).

Flood deposits along Kautz Creek (locality Q; Fig. 4.2) are dominated by unconsolidated, massive, matrix-supported diamicts, with subrounded to well rounded clasts up to 1 m in diameter, supported in a grey-coloured matrix of coarse sand and gravel. Some beds

171 are heterogeneous defined by isolated rafts of massive silty-clay within the clast-rich diamict

(Figs. 4.14A and 4.15). Flood deposits examined along Tahoma Creek (locality R; Fig. 4.2 and

4.14B) and Van Trump drainage basin (locality S; Figs. 4.2, 4.14C) are lithologically similar to those in Kautz Creek. At all localities no glacially striated or shaped clasts could be identified.

4.5 INTERPRETATION

On Mount Rainier, genetic terms such as ‗glacial till‘ or ‗drift‘ have been applied to sediments that are more accurately described as diamictites and interpreted as ‗debrites‘.

Outcrops mapped as ‗till‘ or ‗drift‘ are in fact composed of a mixture of glacial sediment, bedrock rubble, pyroclastic debris and proximal fluvial deposits. They also show well defined bedding parallel to slope (e.g., Figs. 4.5H, I) and normal grading in the most distal beds, which are features typical of deposits emplaced by sediment-gravity flow mechanisms (e.g., Fig.

4.5F, G). In such resedimented diamict deposits, as seen on Mount Rainier, it is typical for clast shape to become more rounded down valley due to the longer transport history of the sediment, as displayed in Figure 4.6 (e.g., Walker, 1975; Nemec and Steel, 1984; Trop et al.,

1999). The terms ‗compound debrites‘ or ‗mixtites‘ are more appropriate terms for these deposits in view of the wide range of lithologies moved by these flows. Large angular boulders typical of rockfall debris carried supraglacially are very common in such deposits and striated clasts are few in number. This is due in part of the nature of the debris being carried on the lower margins of valley glaciers on Mount Rainier which are covered with a very thick mantle of angular supraglacial debris dumped on the ice by rock falls from the high valley walls (Mills, 1978). As a result, ‗glacial‘ sediment deposited by the modern alpine glaciers on

Mount Rainier is dominated by supraglacial rockfall debris that is transported passively on the

172 ice surface (Mills, 1979). Correspondingly, it can be deduced that any subglacial component of sediment (e.g., till) that may have been fingerprinted by striated and shaped clasts is diluted within the total sediment flux moved down valley (see also Eyles, 1979 in relation to Icelandic and European Alpine glaciers). Mills (1979) also recognized the lack of glacially striated and shaped clasts in glacial ‗drift‘ deposits on Mount Rainier but attributed this to the abundance of relatively soft, easily crushed pyroclastic debris. The major factor identified here however, is the mixing of primary glacial sediment (till) with many other sediment types during downslope transport diluting and ‗masking‘ any original glacial sediment and likely also removing striations due to abrasion of clasts by the coarse sandy matrix within the flows. Both Driedger

(1986) and Walder and Driedger (1993) recognized this problem with using the term ‗till‘ and instead, employed the umbrella term ‗glacial drift‘ to encompass slope deposits but the term still exaggerates the influence of ice and disguises the importance of mass flow in the formation of diamicts at Mount Rainier.

As expected, a stronger glacial signal in the form of striated clasts is retained in the continental Vashon Till deposits in the Puget Sound Lowland. Surprisingly, a higher frequency of striated clasts was also observed in deposits of the Osceola Mudflow (lahar), where striated clasts were transported as far as 56 km from their source areas on summit glaciers (Fig.4.12B).

The presence of striated clasts was likely the reason why this deposit was mapped as a ‗till sheet‘ by Willis (1898). In this regard, Futornick (2008) recently noted that several deposits previously mapped as glacial tills in the Puget Sound Lowland are in fact debrites derived from

Mount Rainier (see Pringle et al. 2000, Pringle, 2002; Goldstein et al., 2002). The ability of the

Osceola Mudflow to undercut into glacial, fluvial, bedrock and other volcaniclastic sediment from the volcano has been attributed to the very steep slope (~28o) in the northeast region of

173

Mount Rainier where the Osceola Mudflow flowed down (Mills, 1978). Open voids and vesicles up to 75 cm in diameter are common in the Osceola Mudflow and have been interpreted as ice casts recording the melt of entrained ice blocks (Mills, 1978). Riehle et al.

(1981, p. 8) described active lahars from Mount Redoubt in Alaska that carried numerous broken ice blocks derived from upvalley glacier margins and commented that ice was an important source of water within lahars.

The down valley increase in clast rounding evident in the Osceola Mudflow (Fig.

4.12A) most likely reflects the incorporation of fluvial gravels during flow and the abrasion of clasts due to clast collisions during flow. Data also indicates that deposition of the Osceola

Mudflow was not a simple single event but involved distinct surges and changing flow conditions. Reverse grading and interbedding observed in distal Osceola Mudflow deposits

(locality O; Fig. 4.4) are typical of distal resedimented debris flow deposits in which dilution and the onset of partial turbulence has occurred. Vallance and Scott (1997) suggested that the mudflow was likely emplaced by incremental accretion (separate pulses of deposition) and this is borne out by sedimentological descriptions presented here.

Glacial outburst flood deposits too are dominated by diamicts as a result of the rapid

‗bulking up‘ of floodwaters and the downslope generation of hyperconcentrated flows and debrites. In this case however, it can be deduced that violent clast-on-clast collisions within the body of the flow played a major role in the destruction of striae that may have been present on clasts.

174

4.6 DISCUSSION

Data presented in this study highlight the importance of downslope reworking of glacial and volcanic sediment on Mount Rainier as has been identified at other glaciated volcanoes

(e.g., Mount Redoubt in Alaska; Riehle et al., 1981). Results clearly demonstrate that mass flow is the dominant sedimentary process due to the high-relief and tectonically active setting with frequent volcanic eruptions and abrupt melting of ice, high rain and snowfalls. In short, tills, lahars and outburst flood deposits all record the remobilization and mixing of glacial, volcanic, and fluvial sediment together downslope, resulting in the formation of poorly sorted, debrites.

Many glaciated valleys within the high-relief mountainous North American Cordillera are filled with reworked glacial sediments that were remobilized as debrites during a short- lived phase of heightened slope activity immediately following deglaciation (Eyles et al., 1988;

Eyles and Eyles, 1989). As a result of post-depositional mass flow, primary glacial sediments are redeposited down valley as coarse-grained debrites that exhibit little or no sedimentary evidence of a prior glacial history. Past practice has been to label poorly sorted sediments as

‗till‘ simply because of their poorly sorted outcrop appearance (e.g., Willis, 1898) and a lack of knowledge at that time of alternate depositional processes. The work reported here confirms the findings of ongoing research that indicates that many deposits mapped as glacial tills in the

Puget Sound Lowland are in fact debrites sourced from Mount Rainier (Pringle et al. 2000,

Pringle, 2002; Goldstein et al., 2002; Futornick, 2008).

Glacial sedimentologists have long recognized that tills can be remobilized downslope as debris flows (e.g., ‗flow tills‘: Boulton, 1971, 1978; Lawson, 1979, 1982; Paul and Eyles,

1990; Benn and Evans, 2010). Typically, resedimentation takes place on ice-cored topography

175 underlain by dead ice that is slowly ‗down-wasting‘. This gives rise to continuous readjustments of the sediment cover by mass flow as the topography evolves; topographic depressions formed by more rapid melt fill with debris and become topographic highs as surrounding areas stripped of their debris cover melt down more quickly resulting in

‗topographic inversion‘ (Eyles, 1979). The geographic scale of downslope transport is very limited however and cannot be compared with that on Mount Rainier where debris is moved downslope often for tens of kilometres. ‗Flow tills‘ are moved no more than a few tens of metres downslope before debris is entirely disaggregated by either meltwater or gravity as ice- cored topography ‗down-wastes‘. Long distance transport of glacial sediment within large volume lahars such as on Mount Rainier, involving flow paths several tens of kilometers in length, has seldom been addressed.

Interestingly near Mount Rainier, a thick succession of subaqueously deposited lahars

(diamictites), sandstones and laminated facies, are preserved in the Miocene (~ 12 Ma) lacustrine Ellensburg Formation (Mackin, 1961; Smiley, 1963; Pringle, 2008). They record down valley transport of volcaniclastic sediment along the eastern foothills of ancestral

Cascade volcanoes into a deep lake by lahars. Although no evidence of a glacial contribution to the total sediment flux of this basin has been identified it may provide some insight into the depositional origin of the Konnarock Formation of southwestern Virginia, USA (c. 760-570

Ma; Aleinikoff et al., 1991), which represents one of the Avalonian diamictites linked to the

Gaskiers glaciation. The Konnarock Formation comprises a thick (~ 1 km) succession of subaqueously deposited diamictite, sandstones, and rhythmically laminated mudstones containing ‗outsized‘ clasts, which Miller (1994) interprets as ice-rafted dropstones. Miller

(1994) interprets the succession as being deposited by or near grounded or floating glaciers in a

176 small ice-contact lake close to local glaciers on nearby volcanic centers (Rankin, 1993). A detailed facies analysis of the Ellensburg Formation is beyond the scope of the present study but these rocks are relevant to the understanding of the fate of terrestrial glacial sediment in tectonically active and volcanically influenced pre-Pleistocene glaciated basins, such as in the

Neoproterozoic. The Ellensburg Formation may provide a rare example of reworked glacial and volcanic sediment preserved in a terrestrial volcanic arc setting comparable to the

Konnarock Formation. Further detailed investigation and comparison of the Ellensburg and

Konnarock Formations is clearly warranted.

In the case of the Gaskiers glaciation (c. 580 Ma), ice covers nucleated on volcanic arcs on the outer margin of Gondwana (Avalonian-Cadomian Terranes) and, as demonstrated by the

Gaskiers Formation, were able to reach sea level and influence sedimentation at least in one basin. However, any record of direct terrestrial glaciation (e.g., true tillites) is missing (Eyles and Eyles, 1989; Eyles, 1990). Data collected from Mount Rainier provide an answer by showing that the depositional processes operating in volcanic arc environments are very effective in destroying any terrestrial glacial record on the landward margins of volcanoes; if ice was present any primary glacial sediment would have been reworked downslope during eruptions and meltwater events and subsequently destroyed by the same slope processes at work on Mount Rainier.

4.7 CONCLUSIONS

Volcanic-glacial sedimentary processes on Mount Rainier are dominated by mass flows of glacial and volcaniclastic debris and their downslope mixing and accumulation as debrites

(lahars). Key factors are the high-relief and tectonically active setting with frequent volcanic

177 eruptions and abrupt melting of ice, combined with high rainfall and rapid spring runoff. As a result of the high frequency of mass flow activity on Mount Rainier, primary glacial sediments deposited on its summit and upper slopes are continuously reworked and redeposited down valley as coarse-grained debrites that exhibit little or no sedimentary evidence of a prior glacial history. Therefore, the term ‗debrite‘ is a more appropriate term for describing deposits previously mapped as ‗tills‘ or more generally as ‗glacial drift‘ on Mount Rainier.

Given its tectonic setting, Mount Rainier provides a very good modern depositional analogue for pre-Pleistocene glaciated volcanic settings such as those that occurred during the

Gaskiers glaciation at c. 580 Ma along the Avalonian-Cadomian Terrane (e.g., Gaskiers

Formation). Based on the results of this study it can be deduced that the absence of terrestrial glacial facies (tillites) along the Avalon Terrane can be attributed to the reworking of primary glacial sediment by mass flow processes such as documented here on Mount Rainier. The sedimentary data presented here carry significant implications for identifying and understanding the depositional processes which challenge our ability to recognize terrestrial glacial deposits on the landward margins of ancient glaciated volcanically influenced basins.

178

Figure 4.1 Schematic cross section through the Juan de Fuca Ridge and the Cascadia subduction zone showing location of Mount Rainier within the Cascade Range (modified from Pringle, 2008). Regional map of Mount Rainier is shown in Figure 4.2. Error!

179

B J I A E G F R S Q H D C

Figure 4.2 General schematic of the distribution of glacial drift and outburst flood deposits, bedrock and present-day glaciers on Mount Rainier (modified from Pringle, 2008). Vashon Till study sites (Sites A and B; solid black circles), Evans Creek glacial drift study sites (Sites C-H; vertically-lined circles), Hayden Creek glacial drift study sites (Sites I and J; solid white circles); Outburst flood deposit study sites (Sites Q-S; dotted circles).

180

Figure 4.3 Ice-contact columnar jointed andesite lava flows at Mount Rainier (Burroughs Mountain).

181

O P

L N K M

Figure 4.4 Extent of major lahars generated at Mount Rainier over the past 6000 years (modified from Pringle, 2008) and locations of Osceola Mudflow study sites within the White River valley (Sites K to P; checkered circles).

182

Figure 4.5 (A) Evans Creek glacial drift (locality C; Fig. 4.2) overlain by a thin bed of the Paradise Lahar; contact occurs along dashed line. See Figure 4.9A for corresponding stratigraphic log. (B) Evans Creek glacial drift (locality D; Fig. 4.2).

183

D E

Figure 4.5 (C) Evans Creek glacial drift (locality E; Fig. 4.2). (D) Evans Creek glacial drift (locality F; Fig. 4.2) overlying striated rhyolite.

184

30 cm

Figure 4.5 (E) Enlargement of Evans Creek glacial drift shown in Figure 4.5D (locality F; Fig. 4.2). See Figure 4.9B for corresponding stratigraphic log. (F) Evans Creek glacial drift (locality G; Fig. 4.2) showing normal grading and overlying basalt (contact is marked by dashed line).

185

15 cm

Figure 4.5 (G) Enlargement of Evans Creek glacial drift shown in Figure 4.5F (locality G; Fig. 4.2). (H) Evans Creek glacial drift (locality H; Fig. 4.2) showing well- defined bedding.

186

20 cm

Figure 4.5 (I) Enlargement of Evans Creek glacial drift shown in Figure 4.5H (locality H; Fig. 4.2). (J) Enlargement of Evans Creek glacial drift shown in Figure 4.5H (locality H; Fig. 4.2).

187

D C E G F

Figure 4.6 Scatter diagram generated by plotting the degree of clast roundness in Evans Creek glacial drift deposits (examined at locations shown in Figure 4.2) as a function of elevation. Clast roundness was examined using the visual roundness chart of Krumbein (1941) and the corresponding ‗Roundness Indexes12‘ of Powers (1953) (see Appendix 1). The ‗best-fit‘ regression line and the correlation coefficient (r) indicate that there is a strong relationship between these two variables confirming that there is a strong down valley trend towards increased roundness of clasts in deposits of Evans Creek Drift.

12 See Appendix. 188

Hammer

15 cm

Figure 4.7 (A) Hayden Creek glacial drift (locality I; Fig.4.2). (B) Enlargement of Hayden Creek glacial drift shown in Figure 4.7A (locality I; Fig. 4.2).

189

A B

30 cm

Figure 4.8 (A) Vashon Till deposit in the Ohop Valley in the Puget Sound Lowland (locality A; Fig. 4.2). (B) Enlargement of Vashon Till shown in Figure 4.8A.

190

A) locality F; A) Fig. D; Figs. Fig. 4.2; Fig. 4.2; 4.5B, locality (B)

( .

Stratigraphic logs of key outcrops of Evans Creek Evans of key of logs Drift outcrops Stratigraphic D/E. 4.5

Figure 4.9 Figure

191

), (D) Vashon Till (locality Till (locality Vashon (D) ),

(localityFig. G; 4.2; Figs. 4.5F/G

Stratigraphic logs of key outcrops of (C) Evans Creek Drift Evans Creek (C) of key of logs outcrops Stratigraphic Fig. A; 4.2/ Figs. 4.8A/B).

Figure 4.9 Figure

192

B C

Figure 4.10 (A) Massive Osceola Mudflow deposit (outlined in rectangle) exposed on the floor of Glacier Basin at locality K; Fig. 4.4 (upper limit of inundation). See Figure 4.13A for corresponding stratigraphic log. (B) Striated andesite boulder within the Osceola Mudflow in Glacier Basin (locality K; Fig. 4.4). (C) Striated andesite clast found in the Osceola Mudflow (locality K; Fig. 4.4).

193

1 m

Figure 4.11 (A) Thick and massive succession of the Osceola Mudflow overlying Evans Creek drift (contact indicated by dashed line) exposed along valley walls above the White River (locality L; Fig. 4.4). (B) Massive Osceola Mudflow deposit exposed along Sunrise Road (locality M; Fig. 4.4). 194

35 cm

Figure 4.11 (C) Massive Osceola Mudflow deposit at locality N (Fig. 4.4). See Figure 4.13B for corresponding stratigraphic log. (D) Enlargement of Osceola Mudflow shown in Figure 4.11C (locality N; Fig. 4.4).

195

E F

1 m

Figure 4.11 (E) Osceola Mudflow in the town of Greenwater at locality O (Fig. 4.4) composed of massive and reversely graded diamict beds and interbeds of massive and crudely laminated fine-grained silt-sandstone sediment. See Figure 4.13CA for corresponding stratigraphic log. (F) Enlargement of Osceola Mudflow shown in Figure 4.11E (locality O; Fig. 4.4) showing beds of massive diamict and massive and crudely laminated fine-grained silt-sandstone.

196

75 cm

25 cm

Figure 4.11 (G) Enlargement of fine-grained crudely laminated silty-sandstone bed of the Osceola Mudflow shown in Figure 4.11F (locality O; Fig. 4.4). (H) Enlargement of massive, matrix-supported diamict bed of Osceola Mudflow shown in Figure 4.11F (locality O; Fig. 4.4).

197

5 m

65 cm

2 m

2 m Figure 4.11 (I) Osceola Mudflow at locality O (Fig. 4.4) showing reverse grading defined by upward transition from pebble/cobbles to boulders (along arrow). (J) The most distal deposit of the Osceola Mudflow examined at Mud Mountain Dam (locality P; Fig. 4.4) showing bedding.

198

N L M

P O

B 8 K 7

6 R2= 0.91

4 N M 3

2 Number of Striated Clasts O 1

0 P 0 200 400 600 800 1000 1200 1400 1600 1800 2000 Elevation (MASL)

Figure 4.12 (A) Scatter diagram generated by plotting the degree of clast roundness in Osceola Mudflow deposits examined at locations shown in Figure 4.4 as a function of elevation. The ‗best-fit‘ regression line and the correlation coefficient (r) indicate that there is a strong relationship between these two variables confirming that there is a strong (and predictable) down valley trend towards increased roundness of clasts in deposits of the Osceola Mudflow. (B) Scatter diagram generated by plotting the abundance of striated clasts in Osceola Mudflow deposits examined at locations shown in Figure 4.4 as a function of elevation. The ‗best-fit‘ regression line and the correlation coefficient (r) indicate that there is a strong relationship between these two variables confirming that there is a strong down valley trend towards a reduction in the abundance of striated clasts in deposits of the Osceola Mudflow.

199

C A) locality K; Fig. 4.4; /4.10A, (B) locality N; Fig. N; locality locality /4.10A, 4.4; A) (B) Fig. K;

( .

I. I.

ratigraphic logs of key Mudflow Osceola of key of logs outcrops ratigraphic

St Fig. O; locality 4.4/4.11 (C) 4.11C/D, / 4.4 E and

Figure 4.13 Figure

200

Figure 4.14 (A) Outburst flood deposit exposed in Kautz Creek (locality Q; Fig. 4.2). This deposit records the 1947 flood event derived from the Kautz Glacier. (B) Tahoma Creek outburst flood deposit recording flood events in the 1960‘s (locality R; Fig. 4.2).

201

Figure 4.14 (C) Outburst flood deposit exposed in the Van Trump Creek deposited at circa. 2001(locality S; Fig. 4.2).

202

Fig. Q; 4.2/ cality 4.14A).

- -

Massive,matrix Massive,matrix

supporteddiamictite supporteddiamictite

Massive,matrix

supporteddiamictite

C S F M C M F S C

4 m 4 m 2

3 m 3

5 m 5

1 m 1

Stratigraphic log of Kautz Creek Outburst (lo Kautz deposit flood Creek of log Stratigraphic

Figure 4.15 Figure

203

REFERENCES

Aleinikoff, J.N., Zartman, R.E., Rankin, D.W., Lyttle, P.T., Burton, W.C., & McDowell, R.C. (1991). New U-Pb zircon ages for rhyolite of the Catoctin and Mount Rogers formations: more evidence for two pulses of Iapetus rifting in the central and southern Appalachians. Geological Society of America Abstracts with Programs, 23(1), 2.

Allen, P. A., & Etienne, J.L. (2008). Sedimentary challenge to Snowball Earth. Nature Geoscience, 1, 817.

Arnaud, E., & Eyles, N. (2006). Neoproterozoic environmental change recorded in the Port Askaig Formation, Scotland: Climatic vs. tectonic controls. Sedimentary Geology, 183, 99-124.

Benn, D.I. & Evans D.J.A (2010). Glaciers and Glaciation. 2nd Edition, London, Hodder Education.

Boulton, G.S. (1971). Till genesis and fabric in Svalbard, Spitsbergen. In: R.P. Goldthwait, (Ed.) Till: a symposium. Columbus, OH, Ohio State University Press, 41-72.

Boulton, G. S. (1978). Boulder shapes and grain size distribution of debris as indicators of transport paths through a glacier and till genesis. Sedimentology, 25(6), 773-99.

Boulton, G. S. and Eyles, N. (1979). Sedimentation by valley glaciers: a model and genetic classification. In: C. Schliichter (Ed.) Proceedings of an INQUA symposium on genesis and lithology of Quaternary deposits, Zurich. Moraines and varves: origin, genesis, classification. Rotterdam, A. A. Balkema, 11-23.

Bruckner, M.D. & Anderson, M.M. (1971). Late Precambrian glacial deposits in southeastern Newfoundland, a preliminary note. Geology Association of Canada, Proceedings, 24, 95-102.

Campbell, N.P. (1975). A geologic road log over Chinook, White Pass, and Ellensburg to Yakima highways. Washington Division of Geology and Earth Resources Information Circular, 54, 82.

Carto, S.L., & Eyles, N. (2011). The Gaskiers Formation, eastern Newfoundland, Canada. In: Arnaud, E., Halverson, G. P. & Shields-Zhou, G.A. (Eds.), The Geological Record of Neoproterozoic, Geological Society of London, Memoir, (in press).

Clapman, M. E. & Corsetti, F.A. (2005). Deep valley incision in the terminal Neoproterozoic (Edicaran) Johnnie Formation, eastern California, USA: Tectonically or glacially driven? Precambrian Research, 141, 154-164.

204

Crandell, D.R., & Waldron, H.H. (1956). A recent volcanic mudflow of exceptional dimensions from Mount Rainier, Washington. American Journal of Science, 254, 349-362.

Crandell, D.R. (1963). Paradise debris flow at Mount Rainier. Short papers in geology and hydrology. U.S. Geological Survey Professional Paper, 475-B, 135-139.

Crandell, D.R., & Mullineaux, D.R. (1967). Volcanic hazards at Mount Rainier, Washington. Bulletin: U.S. Geological Survey, 1238, 26 p.

Crandell, D.R. (1969). Surficial geology of Mount Rainier National Park, Washington. Bulletin: U.S. Geological Survey, 1288, 41 p.

Crandell, D.R. (1971). Postglacial lahars from Mount Rainier volcano, Washington. U.S. Geological Survey Professional Paper, 677, 75 p.

Crandell D.R. & Miller, R.D. (1974). Quaternary stratigraphy and extent of glaciation at Mount Rainier region, Washington. U.S. Geological Survey Professional Paper, 847, 59 p.

Crowell, J.C. (1999). Pre-Mesozoic ice ages: their bearing on understanding the climate system. Memoir: Geological Society of America, 192, 106 pp.

Dott, R.H. Jr. (1961). Squantum ―tillite‖, Massachusetts—evidence of glaciation or subaqueous mass movements? Bulletin: Geological Society of America, 72, 1289-1305.

Driedger, C.L. (1986). A visitor‘s guide to Mount Rainier glaciers: Pacific Northwest National Parks and Forests Association, 80 p.

Driedger, C.L., & Kennard, P.M. (1986). Ice volumes on Cascade volcanoes: Mount Rainier, Mount Hood, Three Sisters, and Mount Shasta. U.S. Geological Survey Professional Paper, 1365, 28 p.

Driedger, C.L., & Fountain, A.G. (1989). Glacier outburst floods at Mount Rainier, Washington state, U.S.A. Annals of Glaciology, 13, 51-55.

Eyles, N. (1979) Facies of supraglacial sedimentation on Icelandic and Alpine temperate glaciers. Canadian Journal of Earth Sciences, 16 (7), 1341-1361.

Eyles, C.H., & Eyles, N. (1983). A glaciomarine model for Upper Precambrian diamictites of the Port Askaig Formation, Scotland. Geology, 11, 692-696.

205

Eyles, N. Eyles, C. H. & McCabe, A. M. (1988) Late Pleistocene subaerial debris flow facies of the Bow Valley, near Banff, Canada Rocky Mountains, Sedimentology, 35, 465-480.

Eyles, C.H. & Eyles, N. (1989) The Late Cenozoic White River 'Tillites' of Southern Alaska: Subaerial slope and fan delta deposits in a strike-slip setting. Bulletin: Geological Society of America, 101, 1091-1102.

Eyles, N., & Eyles, C.H. (1989). Glacially-influenced deep-marine sedimentation of the Late Precambrian Gaskiers Formation, Newfoundland, Canada. Sedimentology, 36, 601-620.

Eyles, N. (1990) Late Precambrian 'tillites' of the Avalonian-Cadomian belt; marine debris flows in an active tectonic setting. Palaeogeography, Palaeoclimatology, and Palaeoecology, 79, 73-98.

Eyles, N. (1993). Earth‘s glacial record and its tectonic setting. Earth-Science Reviews, 35, 1 – 248.

Eyles, C.H., & Eyles, N. (2000). Subaqueous mass flow origin for Lower Permian diamicites and associated facies of the Grant Group, Barbwire Terrace, Canning Basin, Western Australia. Sedimentology, 47, 343-356.

Eyles, N., & Januszczak, N. (2004). ‗Zipper-rift‘: A tectonic model for Neoproterozoic glaciations during the break-up of Rodinia after 750 Ma. Earth Science Reviews, 65, 1-73.

Eyles, N., & Januszczak, N. (2007). Syntectonic submarine mass flows of the Neoproterozoic Otavi Group, Namibia: where is evidence of global glaciation? Basin Research 19, 179-188.

Eyles, C.H., Eyles, N. & Grey, K. (2007) Palaeoclimate implications from deep drilling of Neoproterozoic strata in the Officer Basin and Adelaide Rift Complex of Australia: a marine record of wet-based glaciers. Palaeogeography, Palaeoecology, Palaeoclimatology 248, 291- 312

Eyles, N. (2008). Glacio-epochs and the supercontinent cycle after 3.0 Ga: Tectonic boundary conditions for glaciation. Palaeogeography, Palaeoclimatology, Palaeoecology, 258, 89–129.

Fiske, R.V., Hopson, C.A., & Waters, A.C. (1963). Geology of Mount Rainier National Park, Washington. U.S. Geological Survey Professional Paper, 444, 93 p. Frank, V.A. (1985). Hydrothermal processes at Mount Rainier (Ph.D Dissertation): Seattle, University of Washington, 195 p.

Futornick, Z.O. (2008). Sedimentological evidence for a glacial outburst flood and resulting debris flow: Puget Sound Lowland, Washington State.

Goldstein, B. S., Pringle, P. T., & Anderson, N. A. (2002). A late-glacial outburst flood from glacial Lake Carbon, Washington State, USA (abstract). Geological Society of America Abstracts with Programs, 34(5), A-24.

206

Green, N.L. Armstrong, R.L., Harakal, J.E., Souther, J.G., & Read, P.B. (1988). Eruptive history and K-Ar geochronology of the late Cenozoic Garibaldi volcanic belt, southwetsren British Columbia. Bulletin: Geological Society of America, 100, 563-567.

Halverson, G., Hoffman, P.F., Schrag, D.P., Maloof, A.C., & Rice, A.H.N. (2005). Toward a Neoproterozoic composite carbon isotope record. Bulletin: Geological Society of America, 117, 1181–1207.

Hoffman, P.F., Kaufman, A.J., Halverson, G.P., & Schrag, D.P. (1998). A Neoproterozoic Snowball Earth. Science, 281, 1342–1346.

Hoffman, P.F., & Schrag, D.P. (2000). Snowball Earth. Scientific American, 282, 62–75.

Hoffman, P. F., & Schrag D. P. (2002). The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14 (3) 129-155.

Hoffman, P.F., & Li, Z.X. (2009). A palaeogeographic context for Neoproterozoic glaciation Palaeogeography, Palaeoclimatology, Palaeoecology, 277, 158–172.

Jensen, P.A., & Wulff-Pedersen, E. (1996). Glacial or non-glacial origin for the Bigganjargga tillite, Finnmark, northern Norway. Geological Magazine, 133(2), 137-145.

Lawson, D.E. (1979). Sedimentological analysis of the western terminus region of the Matanuska Glacier, Alaska. US Army Cold Regions Research and Engineering Lab, 79-9.

Lawson, D.E. (1982). Mobilization, movement and deposition of active subaerial sediment flows, Matanuska Glacier, Alaska. Geology, 90(3), 279-300.

Lescinsky, D. T., & Sisson, T. W. (1998). Ridge-forming, ice-bounded lava flows at Mount Rainier, Washington. Geology, 26, (4), 351-354.

MacDonald, F.A. Schmitz, M.D., Crowley, J. L., Roots, C.F., Jones, D.S. Maloof, A.C., Strauss, J.V., Cohen, P. A., Johnston, D.T., & Schrag, D.P. (2010). Calibrating the Cryogenian. Science, 327, 5970, 1241-1243.

MacGabhann, B. A. (2005). Snowball Earth, the Dalradian Supergroup, and the Cryogenian- Ediacaran Boundary. 48th Irish Geological Research Meeting Programme and Abstracts, 33.

Mackin, J.H. (1961). A stratigraphic section in the Yakima Basalt and Ellensburg Formation in south-central Washington. Washington Division Mines and Geology, 19, 45 p.

Major, J. J., & Newhall, C. G. (1989). Snow and ice perturbation during historical volcanic eruptions and the formation of lahars and floods. Bulletin of Volcanology, 52, 1–27.

207

Mathews, W. H. (1952). Ice-dammed lavas from Clinker Mountain, southwestern British Columbia. American Journal of Science, 250, 553–565.

McBirney, A.R. (1968). Petrochemistry of Cascade andesite volcanoes, In: Andesite conference guidebook: Oregon Department of Geological and Minerological Industries, Bulletin, 62, 101-107.

Miall, A.D. (1978). Lithofacies types and vertical profile models in braided river deposits: a summary, In: Miall, A.D. (Ed.), Fluvial Sedimentology. Memoir: Canadian Society of Petroleum Geologists, 5, 578- 586.

Mills, H.H. (1978). Some characteristics of glacial sediments on Mount Rainier, Washington. Journal of Sedimentary Petrology, 48 (4), 1345-1356.

Mills, H.H. (1979). Some implications of sediment studies for glacial erosion on Mount Rainier, Washington. Northwest Science, 53, 3, 190-199.

Molen, M. (1990). Diamictites: ice-ages or gravity flows? In: Proceedings of the Second International Conference on Creationsim, R.E. Walsh and C.L. Brooks (Eds.), Creation Science Fellowship, Pittburgh, Pennsylvannia, II, 177-190.

Mullineaux, D.R. (1974). Pumice and other pyroclastic deposits in Mount Rainier National Park, Washington. U.S. Geological Survey Bulletin, 1326, 83 p.

Naranjo, S. J.A., Moreno, R. H., & Banks, N.G. (1993). La erupción del Volcán Hudson en 1991 (46°S), Región XI, Aisén, Chile. Servicio Nacional de Geología y Minería Boletín 44, 50 p.

Nichols, G. (1999). Sedimentology & Stratigraphy. Blackwell Science, Osney Mead, Oxford, 62-75.

Paul, M.A., & Eyles, N. (1990) Constraints on the preservation of diamict facies (melt-out tills) at the margins of stagnant glaciers. Quaternary Science Reviews, 9, 51-69. Pringle, P. (2008). Roadside Geology of Mount Rainier National Park and Vicinity. Washington Division of Geology and Earth Resources, Information Circular 107.

208

Pringle, P. T., Goldstein, B. S., & Anderson, N. R. (2000). Tanwax Creek– Ohop Valley late- glacial flood—Evidence that discharge from an ice dammed lake in the Carbon River Valley was augmented by a temporary landslide dam, Puget Lowland, Washington (abstract). In: Washington Department of Ecology, Washington Hydrologic Society, U.S. Geological Survey, Program and abstracts from the 3rd symposium on the hydrogeology of Washington State, p. 85.

Pringle, P. (2000). Buried forests of Mount Rainier volcano-Evidence for extensive Holocene inundation by lahars in the White, Puyallup, Nisqually, and Cowlitz River Valleys (abstract] Washington Geology, 28(1/2) 28.

Richardson, D. (1968). Glacier outburst floods in the Pacific Northwest, In: Geological Survey Research 1968. U.S. Geological Survey Professional Paper, 600-D, p. D79-D86.

Riehle, J.R., Kienle, J., Emmel, K.S. (1981). Lahars in Crescent River Valley, Lower Cook Islet, Alaska. Geological and Geophysical Report 53, 10p.

Schermerhorn, L. J. G. (1974). Late Precambrian mixtites: Glacial and/or non-glacial? American Journal of Science, 274, 673–824.

Schermerhorn, L. J. G. (1983). Proterozoic glaciation in the light of CO2 depletion in the atmosphere. Memoirs: Geological Society of America, 161, 309– 315.

Scott, K. M., Pringle, P. P., & Vallance, J. W. (1992). Sedimentology, behavior, and hazards of debris flows at Mount Rainier, Washington. U.S. Geological Survey Open-File Report 90-385, 106 p.

Scott, K. M., Vallance, J.W., & Pringle, P. T. (1995). Sedimentology, behavior, and hazards of debris flows at Mount Rainier, Washington. U.S. Geological Survey Professional Paper 1547, 56 p., 1 plate.

Scott, K. M., Macias, J. L., Naranjo, J. A., Rodriguez, S., & McGeehin, J. P. (2001). Catastrophic debris flows transformed from landslides in volcanic terrains—Mobility, hazard assessment, and mitigation strategies. U.S. Geological Survey Professional Paper 1630, 59 p.

Sisson, T.W. (1995). History and Hazards of Mount Rainier, Washington. USGS. Open-File Report 95-642.

Sisson, T.W., & Lanphere, M. A. (1997). The growth of Mount Rainier volcano, Cascade arc, USA: Puerto Vallarta, Mexico, International Association of Volcanology and Chemistry of the Earth‘s Interior General Assembly, p. 5.

209

Smiley, C. J. (1963). The Ellensburg flora of Washington: University of California Publications in Geological Sciences, 35(3), 159-276.

Swanson, D. A., Cameron, K. A., Evarts, R. C., Pringle, P. T., & Vance, J. A. (1989). Cenozoic volcanism in the Cascade Range and Columbia Plateau, southern Washington and northernmost Oregon, Seattle, Washington to Portland, Oregon, July 3–8, 1989—Field trip guidebook T106, 28th International Geological Congress: American Geophysical Union, 60 p.

Vallance, J. W., & Scott, K. M. (1997). The Osceola Mudflow from Mount Rainier— Sedimentology and hazard implications of a huge clay-rich debris flow. Bulletin: Geological Society of America, 109, 2, 143-163.

Vallance, J. W., Breit, G. N., & John, D. A. (2003). Involvement of a hydrothermal system in the catastrophic 5600 B.P. eruption of Mount Rainier (abstract) Geological Society of America Abstracts with Programs, 35, 6, 552-553.

Vance, J. A., Clayton, G. A., Mattinson, J. M., & Naeser, C. W. (1987). Early and middle Cenozoic stratigraphy of the Mount Rainier–Tieton River area, southern Washington Cascades. In: J. E. Schuster (Ed.), Selected papers on the geology of Washington. Washington Division of Geology and Earth Resources Bulletin, 77, 269-290.

Waitt, R.B., & Thorson, R.M. (1983). The Cordilleran Ice Sheet in Washington, Idaho, and Montana, In: H.E. Wright, Jr., (Ed.), Late-Quaternary Environments of the United States, 1: The Late Pleistocene, University of Minnesota Press, 407p.

Walder, J., & Driedger, C. (1993). Glacier debris flows at Mount Rainier, Washington. U.S. Geological Survey, 93-124.

Walker, R.G. (1975). Generalized facies model for resedimented conglomerates of turbidite association. Bulletin: Geological Society of America, 86, 737–748.

Williams, H., & King, A. F. (1979). Trepassey map area, Newfoundland. Canadian Geological Survey Memoirs, 389, 24 pp.

Willis, B. (1898). Drift phenomena of Puget Sound. Bulletin: Geological Society of America, 9, 111-162.

210

Wood, C. & Kienle, A. (1990). Volcanoes of North America: United States and Canada: Cambridge University Press, 354p. 158-160, U.S. Geological Survey Annual Report, 18th, Part 3, 393-436.

Young, G. M. (1995). Are Neoproterozoic glacial deposits preserved on the margins of Laurentia related to the fragmentation of two supercontinents? Geology, 23, 153–156.

211

CHAPTER 5:

VOLCANIC-SEDIMENTARY MASS FLOW FACIES OF THE LESSER ANTILLES ARC AND GRENADA BASIN: MODERN ANALOGUES FOR NEOPROTEROZOIC ‘TILLITES’ OF THE AVALONIAN-CADOMIAN TERRANES

ABSTRACT

Thick (up to 500 m) beds of massive, poorly sorted mixtures of clasts and matrix

(diamictites) are common within many Neoproterozoic deep marine turbidite successions of the Avalonian-Cadomian Terranes, which presently form a tectonostratigraphic belt around the margins of the North Atlantic Ocean (extending from eastern North America to northwestern

Europe). Their ‗till-like‘ appearance first identified almost one hundred years ago, is used uncritically as evidence of a global or near-global glaciation (Gaskiers glaciation: c. 580 Ma, the youngest of four notional ‗Snowball Earth‘ events). This is contrary to results of sedimentological investigations that show that most are non-glacial debrites deposited by mass flow in interlinked volcanic arc basins peripheral to the margin of Gondwana between c. 620-

570 Ma. Diamictites are planar-bedded or broadly-lensate in form, and interbedded with turbidites in very thick (> 7 km) deep marine basin fills forming classic ‗debrite-turbidite associations‘. Syn-depositional volcanic activity is evident and the thickness and sedimentology of these successions reflects the availability of large volumes of ash and other volcaniclastic sediment for remobilization downslope. Diamictites are the product of mixing of mud and conglomerate during submarine slumping. Only in one single case, in the Gaskiers

Formation in Newfoundland, is the presence of local glaciers recorded by ice-rafted striated and faceted dropstone clasts. To better understand the origin of marine Neoproterozoic

212 diamictites within the Avalonian-Cadomian Terranes a detailed facies analysis was conducted on modern subaerial volcaniclastic mass flow deposits exposed on the islands of the Lesser

Antilles Arc (St. Lucia, Martinique, Dominica and Guadeloupe) and their submarine equivalents within the adjacent Grenada Basin using data collected from oceanographic cruises.

The tectonic setting of the Lesser Antilles Arc and the adjacent Grenada Basin is directly analogous to the inferred tectonic setting of the Avalonian-Cadomian Terranes during the

Neoproterozoic (e.g. string of interlinked island arc terranes composed of several active volcanoes and marginal arc basins). The results of this study identify mass flow as the dominant process generating and transporting diamictites into the Grenada Basin, resulting in the accumulation of mega-block deposits and massive diamictites (debrites) within deep marine turbidites. These mass flow facies and their deep marine stratigraphic context are directly comparable to those found within the Avalonian-Cadomian Terranes demonstrating that the ancient diamictites of these terranes can be explained by adopting a model of non- glacial volcanogenic sedimentation typical of modern-day island-arc systems composed of marginal arc basins adjacent to volcanic chains, not severe glacial climates of a Snowball

Earth.

213

5.1 INTRODUCTION AND PURPOSE OF STUDY

The Gaskiers glaciation (c. 580 Ma; Hoffman et al., 1998; Bowring et al., 2003;

MacGabhann, 2005; Knoll et al. 2006) is the youngest of several Neoproterozoic episodes viewed in some quarters as long-lived, extremely cold glaciations with ice covers extending to the equator (‗Snowball Earth Hypothesis‘; Hoffman et al., 1998; Evans, 2000; Hoffman &

Schrag, 2002; Kirschvink, 2002; Schrag et al., 2002; Hoffman, 2005; Macdonald et al., 2010).

In contrast, an alternative argument based on sedimentological, isotopic and faunal work, posits that the glaciations were regional in scale akin to those of the Phanerozoic with wet- based ice masses that were diachronous in their extent and timing (Leather et al., 2002; Condon et al., 2002; Eyles and Januszczak, 2004; Olcott, 2005; Corsetti et al., 2006; Pazos et al., 2008;

Zhao et al., 2009; Zhao and Zheng, 2010).

The global stratotype section for the Gaskiers glaciation outcrops on the Avalon

Peninsula in Newfoundland, Canada (Gaskiers Formation of Williams and King, 1975). It is dominated by marine mass flow deposits, typically debrites and turbidites derived from a volcanic source area (McCartney, 1967; Williams and King, 1975, 1979; Eyles and Eyles,

1989; Eyles, 1990; Eyles and Januszczak, 2004; Carto and Eyles, 2011b). The presence of ice in the volcanic hinterland is indicated solely by dropstones and rare striated and faceted clasts

(Bruckner & Anderson, 1971; Williams & King, 1979; Eyles and Eyles, 1989) indicating that ice was able to reach sea level as floating ice tongues. Glaciation is interpreted to have occurred along the Avalonian-Cadomian Orogenic Belt (also referred to as the peri-

Gondwanan terranes) that was positioned along the outer Gondwanan margin during the Late

Neoproterozoic (Hambrey and Harland, 1981; Hoffman et al; 1998; Bingen et al. 2005;

214

Halverson et al., 2005; Kawai et al., 2008; Hoffman and Li, 2009). Other Neoproterozoic (c.

670-550 Ma) ‗glacial‘ deposits of the Avalonian-Cadomian Terranes, presently scattered around the margin of the North Atlantic Ocean, consist of thick (> 7 km) debrite-turbidite successions and have been classically correlated with the Gaskiers Formation simply on the basis of the till-like character of diamictites (see Hambrey and Harland, 1981 for review).

These include the Squantum Member (‗Squantum Tillite‘ of Sayles, 1919) of the Boston Basin in eastern Massachusetts, the Granville Formation in northern France (Winterer, 1963) and the

Konnarock and the Grandfather Mountain Formations of the Appalachians in eastern USA

(Rankin, 1975; Schwab, 1981; O'Brien et al., 1983; Rast & Skehan, 1983; Rankin, 1993; Miller

1994). Unfortunately there is a scarcity of reliable age data and a lack of substantive evidence for the presence of ice other than the diamictites, which continue to be interpreted as tillites and assumed to be correlative to the Gaskiers glaciation (Kawai et al., 2008; Hoffman and Li,

2009). On the other hand, faunal and sedimentological evidence indicates a deep marine depositional setting for these successions free of any glacial influence dominated by a supply of volcaniclastic sediment. Diamictite facies occur interbedded with thick successions of turbidites derived from the reworking of volcanic ash and are widely identified as debrites produced by the mixing of gravel and mud (Crowell, 1957; Winterer, 1963; Schwab, 1981;

Schermerhorn, 1974; Eyles, 1990, 1993).

A modern tectono-sedimentary analogue for the debrite-turbidite successions of the

Avalonian-Cadomian Terranes of North America and northwestern Europe is the intra-oceanic

Lesser Antilles Arc and the deep water marginal Grenada Basin of the Caribbean. The Grenada

Basin contains 4-7 km of marine volcanogenic sediment comprising turbidites and debrites. On many islands, ‗till-like‘ deposits of subaerial debris flows (lahars) are the dominant

215 sedimentary facies types; similarly gravity flow is the predominant process transporting volcaniclastic sediment down into the Grenada Basin (Sigurdsson et al., 1980; Deplus et al.,

2001; Picard et al., 2006).

This paper presents the results of a field investigation of mass flow facies exposed in lengthy outcrops on four islands of the Lesser Antilles Arc and within the Grenada Basin.

Outcrop data are complemented by results of offshore coring and geophysical surveying by oceanographic cruises (GS-7605, R/V GILLIS, 1976; En-20, R/V ENDEAVOR, 1978;

CARAVAL, R/V L‘Atalante, 2002) in the Grenada Basin. Together, onshore and offshore data provide key insight into the development of debrites along the Lesser Antilles Arc and their transport into the adjacent Grenada Basin. This thesis ends with a comparison of

Neoproterozoic basin fills of the Avalonian-Cadomian Terranes with that of the Lesser Antilles

Arc and the Grenada Basin and discusses the implications of the data amassed from this study for paleoclimate models, particularly the Snowball Earth Hypothesis.

5.2 GEOLOGICAL SETTING AND TECTONIC HISTORY OF THE LESSER ANTILLES ARC SYSTEM

The Lesser Antilles Arc (LAA) and Grenada Basin are part of the Caribbean Plate, which is located between the North and South American, and Nazca and Cocos Plates. The

LAA occurs along the eastern boundary of the Caribbean Plate (Bouysse et al., 1985; Picard et al., 2006) consisting of a 850-km-long intra-oceanic volcanic island arc composed of 30 islands

(archipelago) extending from 12° to 18° N from Grenada to Sombrero (Fig. 5.1A). It is not known whether the Caribbean plate and LAA developed in the late Mesozoic while both were in the Pacific, followed by their eastward migration into their present position between North

216 and South America (‗Pacific Model‘; Pindell & Barrett, 1990; Pindell et al., 2006) or whether they formed in their present day position (‗Inter-America‘ or ‗In-Situ model‘ of Meschede &

Frisch, 1998; James, 2006). Regardless, it is agreed that the LAA has been in its present location since at least the Eocene (~40 Ma), and has been active since then (see Tomblin, 1975;

Lewis and Robinson, 1976; Bouysse et al., 1990; Bouysse and Mascle, 1994).

The LAA splits in two branches north of Martinique (Fig. 5.1A); a western part referred to as the ‗Inner or Recent Arc‘ and an eastern segment known as the ‗Outer or Older arc‘

(Bouysse et al., 1990). The volcanoes of the ‗Outer Arc‘ islands are extinct and covered by thick (up to 500 m) Oligocene–Neogene age limestone. As a result, the islands now form carbonate platforms (the so-called ‗Limestone Caribbees‘ of Mascle and Westercamp, 1983;

Bouysse and Mascle, 1994). The volcanoes of the ‗Inner Arc‘ have been active since 20 Ma, as a result of the slow westward-dipping subduction of the oceanic crust of the western central

Atlantic Ocean under the Caribbean Plate. The islands of the arc are surrounded by narrow shelves that are characterized by shallow depths (10–100 m) (Sigurdsson et al., 1980; Picard et al., 2006).

The LAA is bordered on the west by the deep water back-arc Grenada Basin and on the east by a chain of discrete forearc basins and an accretionary prism (Pinet et al., 1985) (Fig.

5.1A). The size of the Grenada Basin is estimated to be 9.0 x 104 km2. It extends approximately

640 km north-south and 140 km east-west and has a mean water depth of 2 to 4 km (Bouysse,

1988). By comparison, the Grenada Basin is a much larger arc basin than the Boston Basin in eastern Massachusetts, USA, which is estimated to be 750 km2 in size. However, both basins are comprised of turbidite-dominated successions which, at present, are comparable in thickness (up to ~7 km thick).

217

The Grenada Basin developed in a back-arc tectonic regime between the Paleocene to

Eocene when the LAA separated from the now extinct Aves Ridge, located to the west of the

Grenada Basin (Bouysse, 1988). The Aves Ridge represents a former arc from Early

Cretaceous to early Eocene (Minster and Jordan, 1978; Bouysse, 1984). The basin contains a large volume of volcanically derived sediment estimated to be between 4 to 7 km thick overlying basement rocks (Boynton et al., 1979; Picard et al., 2006).

5.3 STUDY AREAS AND METHODS

Detailed facies analyses were conducted on the volcanic-sedimentary clastic deposits on four islands along the LAA: St. Lucia (14°1′N, 60°59′W), Martinique (14o90‘N, 61o10‘W),

Dominica (15°18′N, 61°23′W), and Guadeloupe (16º 15' N, 61º 35' W) (Fig. 5.1B). Andesitic magmas predominate and there is a record of many explosive eruptions. Field study focused on the relatively young (< 1 Ma) subaerial volcanogenic deposits (including pyroclastic flow and surge deposits, debris avalanche and flow deposits, lahars and minor tephra deposits) and uplifted submarine volcaniclastic rocks (including fluvially-reworked conglomerates and sandstones) that are well exposed on all of the study islands (see Sigurdsson et al., 1980; Carey and Sigurdsson, 1989; Bouysse et al., 1990; Deplus et al., 2001; Picard et al., 2006). Field study involved documenting sedimentary structures and features including the texture (grain size, sorting, and clast shape) and fabric of the mass flow deposits, some basic petrography and lithology, large-scale features (geometry and bed contacts, vertical and lateral facies variations) together with syn- and post-depositional deformation structures (e.g., Miall, 1978; Eyles et al.,

218

1983; Nicholas, 1999). A brief description of the geology, geological setting and study areas of each island is provided below.

St. Lucia

The island of St. Lucia, covering an area of 610 km2, is located in the southern part of the LAA (Fig. 5.1A). The known volcanic history of St. Lucia spans over 8 Ma (Le Guen de

Kerneizon et al., 1983) but the last recorded eruptions occurred between 20 to 40 ka. The most recent volcanic activity occurred in the southwestern part of the island in the Qualibou depression which formed sometime between 6 Ma and 290 ka either as a consequence of sector or caldera collapse (Tomblin 1964, 1965; Wohletz et al., 1986; Mattioli et al., 1995).

Fieldwork targeted areas near to the Qualibou structure, which is characterized by large fans of volcanogenic sediments that dip gently seaward toward the Grenada Basin (Briden et al., 1979)

(Fig. 5.1B).

Martinique

The island of Martinique is also located in the southern segment of the LAA covering an area of 1100 km2 and is one of the most volcanically active islands of the arc (Fig. 5.1A)

(Lacroix, 1904; Perret, 1937). It consists of several juxtaposed volcanic centers such as the recently active Mount Pelée and the presently inactive centers of Mont Conil, Morne Jacob and the Pitons du Carbet, in the northern part of the island (Westercamp et al. 1990). The evolution of Mount Pelée volcano has been marked by three major flank collapses (~0.4 Ma, 100-19.5 ka, and 13.5 ka), which systematically destroyed the western side of the volcano (Westercamp and Traineau 1983; Vincent et al., 1989; Le Friant et al., 2003). The products of the ‗collapse‘

219 events at Mount Pelée are well exposed on the western coast of the island (Vincent et al., 1989) and have been traced down into the Grenada Basin (Boudon et al., 1987). As a result, field study centered on the northwestern and western coasts of the island (Fig. 5.1B).

Dominica

Dominica (750 km2) is a small island situated in the central segment of the LAA and has been very active over the past 100 ka (Fig. 5.1A) (Wadge, 1984). This island has seven major volcanic centers all potentially active (Lindsay, 2005). Between ~40 and 20 ka, the island experienced several sector collapses (Carey and Sigurdsson, 1980; Sparks et al., 1980;

Lindsay et al., 2003). The most recent large-scale explosive eruption was around 28 ka from the Trois-Microtrin center in the south (Sigurdsson 1972; Carey and Sigurdsson 1980). Field investigation focused on recent sedimentary deposits consisting of alluvium, conglomeratic debris flows, and debris avalanche and lahar deposits (< 1.8 Ma) exposed along the northwestern, western and southwestern coasts of the island (Sigurdsson et al., 1980; Monjaret,

1985; Bellon, 1988; Sigurdsson and Carey, 1991) (Fig. 5.1B).

Guadeloupe

Guadeloupe is situated in the central region of the LAA (Fig. 5.1A) and is composed of nine islands (Basse-Terre, Grand-Terre, La Désirade, Petite-Terre, Marie-Galante, Terre-de-

Haut, Terre-de-Bas, Saint-Barthélémy, and Saint-Martin) with a total area of 1703 km2

(Westercamp and Tazieff, 1980; Bouysse et al. 1990). The large composite volcano of La

Soufrière in the southern region of the island of Basse-Terre is the only one to have been active in the last 10 ka, experiencing at least 13 collapse events over its life span in the last 140 ka

220

(Blanc, 1983; Carlut et al., 2000; Komorowski et al., 2002, 2005). Field study targeted the volcanically active southern and western coasts of Basse-Terre (Fig. 5.1B).

5.4 DESCRIPTION OF SUBAERIAL MASS FLOW FACIES

Facies analysis of outcrops on the four islands of the LAA described above identifies two principal categories of volcanic-sedimentary deposits: (1) primary pyroclastic facies

(pyroclastic flow and surge deposits, ignimbrites, tephra, and block and ash deposits) derived directly from column collapse or dome collapse, and (2) secondary mass flow facies (typically debris avalanche deposits, lahars, and fluvial sediments) that record the remobilization of primary pyroclastic material down the flanks of the volcanoes as a result of eruptive activity, sector collapse or heavy rainfall. This study focuses on the secondary mass flow facies, although both groups are commonly interbedded. Two categories of subaerial secondary mass flow facies were examined during field study. These include: (1) debris avalanche deposits and

(2) lahar deposits.

5.4.1 Debris avalanche deposits

On each of the four islands investigated in this study, a very substantial component of outcropping sediment consists of debris avalanche deposits from large landslides. These deposits are exceptionally coarse grained consisting of large blocks of volcanic rock (from less than 1 m in diameter to several hundred metres; the ‗megablocks‘ of Mimura et al., 1982; Ui,

221

1983). They record large-scale slope failure or the sector collapse of the steep-sided volcanic cones (Crandel et al., 1984; Glicken et al., 1989; Bouysse et al., 1990; Glicken, 1991; Deplus et al., 2001; Picard et al., 2006).

Both matrix- and block-supported debris avalanche deposits are present, but the latter facies dominate. They consist of massive and unstructured beds measuring from several metres to up to 500 m thick. These show prominent clusters of boulders and blocks; matrix-supported facies supported by a coarse sand-gravel matrix derived by crushing of volcanic lithologies, dominate in more distal areas near the coastline (Figs. 5.2A, B and 5.6A, B). Both facies types are composed of weathered, hydrothermally-altered and fractured lavas and are angular to subangular in shape. These facies commonly display a fractured ‗jigsaw-like‘ surface pattern where fragments of boulders and megablocks can be visually refitted across fractures (Ui,

1983). Clast lithology is generally homogeneous within a single unit indicating a single source landslide area, although some matrix-supported units are multi-lithic in composition comprised of a mixture of basaltic andesite, andesite, and rhyolite clasts. In plane view, both block-and matrix-supported deposits have a markedly lobate geometry with a distinctive coarse blocky and hummocky surface topography in the form of numerous mounds formed by debris draped over large blocks. These deposits are commonly found along valleys, and appear to be restricted to the western flanks and platforms of the islands, with many occurring downslope of volcanic collapse structures (e.g., the Qualibou structure on St. Lucia and Mount Pelée on

Martinique). Block facies are commonly interbedded with primary pyroclastic deposits often dominated by pumice and scoria indicating repeated landsliding and eruptive activity.

The volume of debris avalanche deposits from these islands is impressive. Over the last

100 ka more than 47 volcanic flank collapse events with an estimated total volume of about

222

110 km3 have occurred from the active volcanoes in the LAA, predominantly in the southern segment part (Guadeloupe to St. Lucia; Bouden et al., 1984; Deplus et al., 2001). Slope failure leading to sector collapse may be caused by factors such as oversteepening in the summit region of the volcano, the emplacement of shallow dikes and subsequent horizontal extension, and large earthquakes (Guest et al., 1984; Siebert, 1984). The observation that much debris has been hydrothermally altered suggests that large-scale weakening of rock mass strength is also a factor. The extreme angularity of the blocks results from collisional breaking of blocks during transport (Cas and Wright, 1987) and the relatively short transport distance, although volcanic debris avalanche deposits are known to be very mobile with long run-out distances for a given height of fall (Siebert, 1984). Matrix-supported block facies, found downslope along the coast suggests that some avalanche flows may outdistance others by being able to generate matrix by fracturing and milling of disintegrating blocks. The exclusivity of debris avalanche deposits to the western slopes of the islands of the arc are interpreted to be due in large part to the steepness of the western slopes (average 9o) compared to the eastern slopes (average 1o) (Le

Friant, 2001).

5.4.2 Lahar deposits

The results of this study confirm previous reports that the dominant sediment type on

LAA islands are diamictites deposited by lahars (Sigurdsson et al., 1980). These record the rapid downslope remobilization and mixing of fresh, unconsolidated pyroclastic ejecta, weathered volcanic rock, water, soil and other debris during or after explosive volcanic eruptions due to seismic disturbances and rapid ground inflation preceding volcanism

223

(Sigurdsson et al., 1980; Voight et al., 1981, 1983; Fisher and Schmincke, 1984; Lavigne,

2000). Heavy rainfall events and phreatomagmatic eruptions are important triggers for downslope mass flow (e.g., Sigurdsson et al., 1980; Smith and Roobol, 1990; Fisher and Smith,

1991).

On the four islands, lahar deposits typically consist of unlithified or partially lithified massive matrix- and clast-supported diamictite facies, as well as rare stratified and graded units. In general, given that lahars are formed by the reworking and mixing of other sediment from several sources in the vent environment, the composition of all diamictite deposits on these islands consists of an array of multi-coloured heterolithic volcaniclastic clasts consisting of andesite, dacite, basalt, basaltic andesite and rhyolite, supported by a fine-grained brown or grey matrix of sand-to silt-sized pyroclastic (ash, crystals and glass shards) and lithic material that can amount to as much as 70 vol. % of the total rock volume in matrix-supported units and

< 20 vol. % of the total rock in clast-supported units (Figs. 5.3A-C and 5.6C). The relative abundance of each clast lithology is highly variable from one island to another. Some clast- supported diamictites were also observed to be monolithic in composition, consisting of mostly andesitic or basaltic clasts, but these units are rare. Overall, clasts typically range in size from a few centimeters to a metre in diameter. Clast shape is also highly variable but dominated by subrounded to rounded shapes, and lack any preferred orientation. It is common to find diamictites associated with blocky debris avalanche deposits and pyroclastic flow and surge deposits; in both cases the boundaries between facies are sharp and planar. Some diamictite units are overlain by thin beds (<30 cm thick) of fine-grained massive to crudely laminated beds of the same lithology as the matrix material of the underlying diamictite (Figs. 5.3D and

5.6C). Fragments of non-carbonized wood can be identified in some deposits.

224

Some beds of matrix-supported diamictite record limited mixing of source sediments such that original components can still be recognized. These deposits show isolated rafts of fine-grained material ‗floating‘ within more homogeneous diamictite (Fig. 5.3E). These

‗heterogeneous‘ facies record incomplete mixing during flow. Large clast clusters are also common suggesting reworking of blocky debris avalanche deposits by debris flow processes. It is also common to find matrix-supported diamictite occurring conformably with clast- supported, massive diamictite units within the same bed or transforming laterally downslope into massive clast-supported conglomerates. A crudely developed reverse grading is locally present in some beds (Fig. 5.3F). Overall, deposits exhibit well-defined bedding, range in thickness from a metre to several metres and occur as tabular beds or irregular sheets and mounds in outcrop, in most cases displaying a locally ropy surface morphology. Most beds extend laterally from tens of metres to a hundred metres. The lack of clast angularity in both matrix- and clast-supported diamictites is indicative of the incorporation of pre-existing fluvial sediment as lahars move down slope. The lack of grading, random clast orientation and high matrix content in matrix-supported diamictites suggest the deposits result from laminar, cohesive flows (Nemec and Steel, 1984; Lowe and Guy, 2000; Mulder and Alexander, 2001;

Shanmugam, 2000, 2002). Rare grading within some lahars (diamictites) indicates the onset of late stage turbulence within flows (e.g., Hiscott, 1994; Sohn et al., 1997). Lahars are easily differentiated from primary pyroclastic flows by their composition; pyroclastic flows occur as mixtures of monolithic and juvenile (angular, blocky and jagged) clasts of volcanic rock, scoria and/or pumice, including volcanic bombs and blocks supported in a matrix of ash (clay- to sand-sized particles of pulverized rock, crystals and pumice fragments). They also lack any

225 non-volcanogenic components, display poor to no bedding and are typically overlain and underlain by pyroclastic surge deposits.

5.5 OFFSHORE MASS FLOW FACIES IN THE GRENADA BASIN

Deep-sea coring and geophysical seismic surveys of submarine deposits in the Grenada

Basin have shown that mass flows generated on-land are a major source for sediment delivered to the basin (Sigurdsson et al., 1980; Bounden et al., 1984; Carey and Sigurdsson, 1987;

Deplus et al., 2001; Picard et al., 2006). Large landslides along steep coasts carry large volumes of terrestrial debris downslope directly into deep water where proximal debris avalanche deposits evolve downslope into lahars and turbidites (see Fisher, 1971; Picard et al.,

2006).

Over the last several decades, more than 200 sediment cores have been recovered from the Grenada Basin (cruises GS-7605, R/V Gillis, 1976; En-20, R/V Endeavor, 1978;

CARAVAL, R/V L‘Atalante, 2002; Roobol and Smith, 1980; Carey and Sigurdsson, 1980;

Sigurdsson et al., 1980; Bouysse et al. 1990; Deplus et al 2001; Picard et al. 2006). Based on this data it has been estimated that over the last 100 ka some 85 km3 of volcanic material has been erupted from the five active islands of St. Lucia, Dominica, Martinique, Guadeloupe and

St. Vincent and 442 km3 of volcanic-sedimentary material from the arc has accumulated in the

Grenada Basin; a small volume of terrigeneous sediment is also transported into the Grenada

Basin by rivers draining the South American continent (Sigurdsson et al., 1980). It is estimated that over the last 100 ka only 20% of the volcanogenic sediment produced at the arc remained on the volcanic islands; the vast majority has been transported into the Grenada Basin by

226 slumps, debris flows and turbidity currents (Sigurdsson et al., 1980). The thickness of the basin fill is estimated to measure between 4 to 7 km (Speed et al. 1993) but there is a marked asymmetry in the distribution of the type and abundance of marine volcanogenic sediments around the arc. The thickest accumulation occurs on the west side of the arc in the Grenada

Basin (debris flows and turbidites make up about 98% and air-fall tephra layers less than 2% of all volcanogenic sediment in the basin). These proportions reflect relatively steep (average of

9o) slopes on the western side of the arc and the dominant westerly direction of wind

(Sigurdsson et al. 1980) (Figs. 5.4 and 5.5). The volcaniclastic component in the sediments of the Grenada Basin ranges from 20 to 90% (Sigurdsson et al., 1980; Picard et al., 2006). The asymmetry in the distribution of the type and abundance of marine sedimentation around the arc system is typical of arc/basin systems (Karig and Moore, 1975; Carey and Sigurdsson,

1984).

Multi-beam echo-sound surveys identify fan-shaped, large-volume ―chaotic deposits‖ with hummocky surfaces radiating from the western upper submarine slopes of Dominica, St.

Lucia, Martinique and Montserrat extending deep into the Grenada Basin (Fig. 5.1B) (Le

Friant, 2001; Deplus et al., 2001; Le Friant et al., 2002; Le Friant et al., 2003). These have been recognized as debris avalanche deposits composed of coarse blocks that measure up to more than 100 m in size (see also Lipman et al., 1988; Watts et al., 1995; Urgeles et al., 1999;

Le Friant, 2001; Deplus et al., 2001; Le Friant et al., 2002; Le Friant et al., 2003). Deposits on the submarine slopes of Martinique and St. Lucia occur at the mouth of flat-floored canyons bound by steep linear walls 3-5 km wide and 13-20 km in length. These channels terminate upslope at sites of sector collapse on Mount Pelée on Martinique and the Qualibou structure on

St. Lucia (Deplus et al., 2001). Debris avalanche deposits cover a large area of the basin floor;

227 approximately 3500 km2 off Dominica, 800 km2 off Martinique and 2000 km2 off St. Lucia, with run-out distances for the flows reaching about 90 km, 60 km and 75 km, respectively.

Overall, debris avalanche deposits have an accumulated coverage area in the basin of at least

6300 km3 (>30% of the basin floor between Dominica and St. Lucia; Roobol et al., 1983;

Vincent et al., 1989; Mattioli et al., 1995). Based on known sedimentation rates of the Grenada

Basin (4-20 cm/ka; Sigurdsson et al., 1980) the age of these debris avalanche deposits is estimated to be between < 100 and 200 ka.

Another major component of the strata found in the Grenada Basin are thick (0.5 m to more than 5 m thick) debrite (diamictite) beds (Sigurdsson et al., 1980). Diamictites are composed of poorly sorted mixtures of subrounded to subangular lithic and pyroclastic clasts

(pebble- to boulder-sized clasts) set in a sand-, silt- and clay-sized matrix composed of lithic fragments, crystals, and glass (75 vol. % of deposit), with secondary components (25 vol. %) such as silty-clay rip-up clasts (Carey and Sigurdsson, 1980; Sigurdsson et al., 1980; Reid et al.

1996; Picard et al. 2006). Diamictites (debrites) are commonly interbedded with beds of hemipelagic grey-brown coloured, silty-clays (Carey and Sigurdsson, 1987) (Fig. 5.7).

Typically the debris flow deposits are described as having a ‗chaotic‘ appearance and being interbedded with silty-clay turbidites (5-10 cm thick) (see below) (Carey and Sigurdsson,

1987). The main source of debris flow is from subaerial debrites from Guadeloupe, Dominica,

Martinique, St. Lucia and St. Vincent (Sigurdsson et al., 1980; Carey and Sigurdsson, 1984;

Reid et al. 1996; Picard et al. 2006).

The predominance of debrites west of the arc is attributed to steep slopes along the western arc flanks leading from the volcanoes which allows terrestrial flows to rapidly pass through the critical region at the coastline where most flows disintegrate due to hydromagmatic

228 explosions that occur from the interaction of hot flow and sea water (Sigurdsson et al., 1980;

Picard et al., 2006). On the east side of the arc, where slopes are less steep (1-2°) flows are prevented from entering the basin because their low flow velocity prevents them from crossing the coastal zones (Sigurdsson et al., 1980). The steepness of the slopes on the western side of the arc may be attributable to the fact that most of the volcanic centers also occur on the western side and typically the submarine slopes of arc-related volcanoes are very steep (Fisher and Smith, 1995). Offshore debrites may also represent the distal equivalents of larger submarine debris avalanches moving downslope into deep water from onshore sector collapses that carry parts of the coastal zone with them. Geophysical data shows that submarine debris avalanche deposits transition into well-bedded ―chaotic‖ horizons. As a result it is proposed that upon entry into the basin the debris avalanches erode and disturb the substratum and likely incorporate a large amount of fine-grained wet marine sediment during transport and, as interpreted by Deplus et al. (2001), transform into ―chaotic‖ debris flow units before attenuating on the basin floor.

Volcaniclastic turbidites are the most abundant lithofacies present in the Grenada Basin

(Sigurdsson et al., 1980; Picard et al., 2006). Piston cores from the eastern half of the Grenada

Basin, near the break in slope, are dominated by massive to normally graded volcanic silt and sandy silt turbidites, interbedded with fine hemipelagic sediment. Turbidites range from 5 to

240 cm in thickness and show sharp bottom and top contacts (Sigurdsson et al., 1980; Carey and Sigurdsson, 1984). Turbidites are composed of abundant sand- and silt-sized fragments of crystals, volcanic glass, detrital feldspar, calcareous microfossils, and minor clay particles

(25%). Some show highly convoluted laminations and many features of fluidized sediment flow attributed to the slumping of littoral sand downslope as a result of earthquakes. It is

229 common for these beds to be interbedded with numerous olive-gray ‗ash turbidites‘ dominated by crystals and volcanic lithic fragments, which are normally graded, forming couplets with hemipelagic mudstone (Picard et al., 2006). Hemipelagic sediment is derived from terrigeneous material transported into the Grenada Basin by fluvial and aeolian processes (e.g., Carey and

Sigurdsson, 1980; Sparks et al., 1980; Carey and Sigurdsson, 1984). Hemipelagic intervals are thickest towards the central part of the basin and near the Aves Ridge due to a decreasing influence of volcaniclastic input. Sand-rich turbidites are largely absent from the more distal western half of the Grenada Basin reflecting the formation of localized submarine fans at the base of the steep western slopes of the arc (Garcia 1996; Picard et al. 2006).

Primary ash-fall deposits are also present and constitute some 2 % of the total basin fill.

They occur as thin (0.5-10 cm thick) layers consisting of fine-grained (< 250 um) glassy shards, variable proportions of crystals (plagioclase, amphibole and quartz), unaltered lithic fragments, and minor amounts (< 5%) of non-volcanic sediment (carbonate and clay). Both massive and laminated layers are recognized with some showing normal grading resulting from differential settling through the water column following fall-out from an eruption cloud (Carey and Sigurdsson, 1980; 1984). Although, ash-fallout layers are by volume the least abundant deposit-type in the basin, they are the most widely distributed facies in the Grenada Basin

(Carey and Sigurdsson, 1984; Picard et al., 2006).

Over the last 100 ka, it is estimated that 527 km3 of volcanic sediment has been produced from the volcanoes of the Lesser Antilles Arc. Based on deep-sea coring of the

Grenada Basin over 318 km3 of this sediment has accumulated in the basin of which volcaniclastic sediment (debris flow and avalanche deposits, lahars and turbidites) constitutes

313 km3 of the total volume and dispersed ash is estimated to contribute 5 km3 to the total

230 volume. The remaining sediment has been preserved on-land (85 km3) and within the Atlantic

Ocean (124 km3) (Sigurdsson et al., 1980). Sedimentation in the Grenada Basin has resulted in the formation of several coalescing volcaniclastic aprons, although it has been noted that the basin fill exhibits some features suggestive of submarine fan development (e.g., grain size reduction basinwards) (Karig and Moore, 1975; Sigurdsson et al., 1980; Carey and Sigurdsson,

1987; Picard et al., 2006).

5.6 IMPLICATIONS FOR UNDERSTANDING NEOPROTEROZOIC PALEOCLIMATE DURING THE GASKIERS GLACIATION AT C. 580 MA

In this section the results of the investigation of the modern mass flow facies of the

LAA and the Grenada Basin are used to improve the understanding of the much-debated paleoclimate during the Neoproterozoic Gaskiers glaciation.

As related above, the Gaskiers glaciation at c. 580 Ma has been classically traced along the Avalonian-Cadomian Orogenic Belt now found scattered around the margins of the North

Atlantic Ocean as a consequence of the breakup of Pangea. Rocks of this Orogenic Belt extend from eastern North America to northwestern Europe. The Orogenic Belt was initiated near the margin of Gondwana at c. 1.2 to 1.0 Ga with wide-scale magmatic activity commencing at c.

640-600 Ma when the orogenic belt began subducting beneath the Gondwanan margin. This resulted in the formation of several magmatic arcs and arc-related basins (Murphy and Nance,

1989; Pe-Piper and Murphy 1989; Pe-Piper and Piper 1989; Smith and Socci 1990; O‘Brien et al. 1996; Murphy et al. 1999). A direct comparison can be made between the tectono- sedimentary setting and deep marine mass flow facies of the Avalonian-Cadomian Belt and that of the LAA and the Grenada Basin. A similar arc-type basin setting close to active

231 volcanoes has been inferred for the volcano-sedimentary successions of the Avalonian-

Cadomian Belt (O‘Brien et al., 1983; Nance et al., 1991; Nance and Murphy, 1994; Murphy et al., 200; Dörr et al., 2004), which are also composed of a considerable thickness of volcanically influenced sediment ranging from 4 to 9 km. Furthermore, Avalonian-Cadomian basin fills are also replete with mass-flow-emplaced facies rich in arc-derived volcanic material, ranging from very coarse and blocky debris avalanche deposits, and massive conglomerate and diamictite debrites interbedded with volcaniclastic turbidites, indicating a deep marine volcanically influenced settings and rapid deposition. Such successions occur in

Newfoundland (Gaskiers Formation; c. 580 Ma; Bruckner, 1977; Williams and King, 1975;

Eyles and Eyles, 1989; Narbonne, 2004), the Boston Basin of New England (the Boston Bay

Group c. 595-570 Ma; Thompson and Bowring, 2000), the Appalachians of North America

(Konnarock and the Grandfather Mountain Formations; c. 760-570 Ma; Rankin, 1993; Miller,

1994), and northern France (Granville Formation c. 550-584 4 Ma of Brittany; Winterer,

1963; Schwab, 1981).

Diamictites of the Boston Basin are interpreted in some quarters as ‗glacial tillites‘

(e.g., Squantum Tillite) deposited during the Gaskiers glaciation (Thompson and Bowring,

2000; Narbonne and Gehling, 2003; Bingen et al. 2005; Halverson et al., 2005; Kawai et al.,

2008; Hoffman and Li, 2009). This interpretation is long standing simply because these facies are poorly sorted and resemble modern tills, despite their deep marine origin (see Rankin,

1993, p. 19) and great thickness (500 m thick in some basins; see Miller, 1994). Further evidence of a ‗glacial‘ origin was based on the interbedding of massive diamictites with laminated sediments interpreted as seasonally-deposited rhythmites (e.g., glaciolacustrine varvites; Sayles, 1914, 1919; Emerson, 1917; Billings, 1929; Billings et al., 1939; Caldwell,

232

1964; Cameron, 1976). In respect to the Boston Basin, the author has argued against any glacial influence on sedimentation, instead identifying the ‗tillites‘ of Sayles (1919) and others

(Sayles and LaForge, 1910; Emerson, 1917; Caldwell, 1964; Cameron and Jeanne, 1976) as debrites produced by downslope mixing of conglomerate and mud (Chapter 2 of this thesis), though recently Passchier and Erukanure (2010) make the case for geographically-limited wet- based glaciation in an open marine environment on the basis of ‗dropstone‘ horizons and geochemical data indicating significant exposure of land surfaces during diamictite deposition

(see Chapter 3 of this thesis for an alternative explanation). The so-called ‗glaciolacustrine rhythmites‘ of the Boston Basin (the laminated argillite facies of the ‗Cambridge Formation‘ identified by Sayles, 1914 and Caldwell, 1964) are marine turbidites, containing deep marine

Ediacaran fauna (Dott, 1961; Schermerhorn, 1974; Bailey, 1987; Thompson and Bowring,

2000; Bailey and Bland, 2000; Bailey, 2005).

In northern France, Crowell (1957), Winterer (1963) and Schwab (1981) have amply demonstrated a deep marine debris flow origin for the diamictites in Brittany. The diamictites and turbidites of the poorly-dated Konnarock Formation in southwestern Virginia have also been investigated and have been interpreted as records of a glacially influenced ice-contact lake setting related to repeated episodes of alpine glacial advance and retreat (Miller, 1994) but yet again, there is no unequivocal sedimentological evidence of glaciation (Schwab, 1981;

Eyles and Eyles, 1989) other than the supposed ‗till-like‘ character of diamictites. In reality, a definite glacial influence on deep marine volcanic arc sedimentation has only been confirmed for the Gaskiers Formation in Newfoundland, Canada, on the basis of striated and faceted clasts and dropstones in turbidites (Eyles and Eyles, 1989; Eyles, 1990; 1993, 2008).

233

It is important to note that mega-blocks of sedimentary and volcanic rocks comparable to those of the Grenada Basin are interbedded with diamictites and turbidites within the

Neoproterozoic (c. 595 to 630 Ma) Avalonian Terrane of southeastern New England (Bell and

Alvord, 1976; Bailey et al., 1989) notably the Westboro Formation of northeastern

Massachusetts (Bell and Alvord, 1976), the Newport Neck Formation of southeastern Rhode

Island (Rast and Skehan, 1981; Webster, 1986) and the Blackstone Group of northeastern

Rhode Island (Quinn et al., 1949). For example, in the case of the Westboro Formation, the stratigraphic succession is extensively slump folded (Bell and Alvord, 1976) and mega-blocks are as much as 1 km in diameter. Beds of poorly sorted diamictite composed of intraclasts and varying amounts of quartzite and carbonate fragments (described as ‗mudstone slump deposits‘; Bailey et al., 1989) are common. The composition of megablocks within southeastern New England suggests a source along the margin of the West African or

Amazonian Craton (Bailey et al., 1989) (c. 650-570 Ma) (Skehan and Rast, 1983; Zartman and

Naylor, 1984). These three formations result from the development of large slumps and debris flows on the upper parts of a steep, rifted submarine slope of a large back-arc basin system

(‗Newport-Westboro-Blackstone Basin‘ of Bailey et al., 1989) similar to the Grenada Basin.

Significantly, as revealed from the results of this study, arc-related basins produce enormous volumes of volcaniclastic detritus, which is reworked downslope as mega-block- diamictite-turbidite successions (Karig and Moore, 1975; Glicken, 1986; Siebert et al., 1987;

Carey and Sigurdsson, 1987). The similarity in deposits and overall stratigraphy between the modern-day LAA and flanking deep marine Grenada Basin with the ancient arc-related basins of the Avalonian-Cadomian Terranes is impressive. By comparison with the mass flow facies of the LAA and the Grenada Basin the thick diamictite-turbidite successions of the Avalonian-

234

Cadomian Terranes can be viewed as non-glacial debrites that reflect a volcanic source area of abundant sediment, steep basin slopes and extremely active tectonism and volcanism.

Glaciation at this time is only confirmed from the type section in Newfoundland,

Canada (Gaskiers Formation), suggesting that glaciers were rare and likely restricted to the summits of the highest volcanic cones along the ancient volcanoes of the Avalonian-Cadomian

Terranes, as previously argued by Eyles and Januszczak (2004). The results presented here do not support the concept of a severe globally widespread Snowball Earth at c. 580 Ma.

5.7 CONCLUSIONS

Thick (up to 500 m) beds of massive, poorly sorted admixtures of clasts and matrix

(‗diamictites‘) within the Neoproterozoic Avalonian-Cadomian Terranes, now widely dispersed around the margins of the North Atlantic Ocean (from eastern North America to northwestern Europe), are classically interpreted as tillites and are linked in some quarters to a putative Snowball Earth event (Gaskiers glaciation) at c. 580 Ma. However, the diamictites of this orogen are conformably interbedded with turbidites in very thick (> 7 km) deep marine basin fills. Only in the Gaskiers Formation in Newfoundland can ice-rafted, faceted and striated clasts be identified indicating glacial ice in the volcanic hinterland. The sedimentology and great thickness of these successions (up to 7 km) reflects the availability of large volumes of ash and other volcaniclastic sediment and its remobilization downslope as debris flows and turbidity currents. The modern-day LAA in the Caribbean is an appropriate depositional and tectonic analogue for the diamictites of the Neoproterozoic Avalonian-Cadomian Belt. Facies data collected from four islands (St. Lucia, Martinique, Dominica and Guadeloupe) combined

235 with offshore data from oceanographic cruises underscore the importance of debris flow processes in the accumulation of diamictites (debrites) within deep marine turbidites of arc- related basins. These data negate a Snowball Earth condition during Gaskiers glaciation at c.

580 Ma and emphasize the need for detailed sedimentological assessments of Neoproterozoic

‗tillites‘ as a necessary pre-condition for paleoclimate modeling.

236

Lesser Antilles Arc

Figure 5.1 (A) General schematic of the geological components (indicated by

shading) of Lesser Antilles island arc system located at the eastern margin of the Caribbean Plate (modified from Picard et al., 2006).

237

Malendure

Grand Anse Beach Wotten Waven Springs

Rosalie Point

Grand Caille Point

Choisel Beach

Figure 5.1 (B) Bathymetry of the eastern and western sides of the Lesser Antilles Arc showing the location of study sites on each island, the lateral extent of different debris avalanche deposits on the sea floor of the Grenada Basin, and the main active volcanoes on each island (modified from Bounden et

al., 1992). The locations of key outcrops referenced in Figures 5.2A, B and 5.3A-F are indicated on this figure. 238

Figure 5.2 (A) Block-supported debris avalanche deposit (Wotten Waven Sulphur Springs, Dominica; Fig. 5.1B). See Figure 5.6A for corresponding stratigraphic log. (B) Block-supported debris avalanche deposit overlying a matrix-supported debris avalanche deposit (Choisel Beach, St. Lucia; Fig. 5.1B). See Figure 5.6B for corresponding stratigraphic log.

239

A B

C D

E F

Figure 5.3 (A) Poorly sorted massive lahar deposit (Wotten Waven Sulphur Springs, Dominica; Fig. 5.1B). (B) Poorly sorted massive lahar deposit (Malendure, Basse Terre of Guadeloupe; Fig. 5.1B). (C) Poorly sorted massive lahar deposit (Grande Anse Beach, Basse Terre of Guadeloupe; Fig. 5.1B). See Figure 5.6C for corresponding stratigraphic log. (D) Moderately sorted massive lahar deposit overlain by a thin bed of crudely laminated fine- grained clay-silt-sand (Grande Anse Beach, Basse Terre of Guadeloupe; Fig. 5.1B). (E) Lahar deposit showing incomplete mixing of finer grained matrix material with coarser grained sediment (Rosalie Point, Dominica; Fig. 5.1B). (F) Massive, matrix-supported lahar that transitions into a clast-supported, reversely graded lahar (contact between these units is indicated by dashed line) (Grande Caille Point, St. Lucia; Fig. 5.1B).

240

Figure 5.4 Schematic illustration of the asymmetric distribution of volcanogenic deposits east and west of the Lesser Antilles Arc and the factors chiefly responsible for this dispersal pattern. The figure portrays an eruption plume that is transported to the east by the prevailing westerlies, resulting in air-ash layers in the Atlantic Ocean east of the arc. Contemporaneous debris flows enter the sea west of the arc where the slopes are steep (modified from Sigurdsson et al. 1980).

241

Figure 5.5 Schematic east-west cross-section of the Lesser Antilles arc system showing the distribution of volcanogenic sediments in and around the active arc (modified from Sigurdsson et al. 1980).

242

Beach, St. Lucia (Figs. 5.1B/ Fig. 5.2B), and (C) Grande Anse Beach, Basse Terre Anse Basse Beach, Grande (C) and 5.1B/ (Figs. St. Lucia Fig. 5.2B), Beach,

D). D). -

Stratigraphic logs of subaerial mass flow facies from the Lesser Antilles Arc. (A) Wotten Waven Sulphur Waven Antilles (A) Arc. flow mass Wotten from Springs, Lesser the facies subaerial of logs Stratigraphic Choisel 5.1B), (Fig. (B) Dominica 5.1B, (Figs. 5.3C Guadeloupe of

5.6 Figure

243

Figure 5.7 Lithology of sediment core retrieved from the deepest part of the Grenada Basin based on deep-sea core named ―CAR-DOM-1‖ (CARAVAL cruise, 2002) located 120 km west of the island of Martinique. Diagram shows interbedding of volcaniclastic ash layers, hemipelagic sediments and thick diamictite deposits (modified from Picard et al., 2006). 244

REFERENCES

Bailey, R.H., Skehan, J.W., Dreier, R, B. & Webster, M.J. (1989). Olistostromes of the Avalonian terrane of southeastern New England. Geological Society of America, Special Paper, 228, 93-112.

Bailey, R.H., & Bland, B.H. (2000). Ediacaran fossils from the Neoproterozoic Boston Bay Group, Boston area, Massachusetts. Geological Society of America Abstracts with Programs, 32, A11.

Bailey, R. H. (2005). Stratigraphy and Depositional Setting of Neoproterozoic Strata on Southern Boston Harbor Islands, Boston Bay Group, Massachusetts. Geological Society of America, Northeast Section (Abstract), 27(7), 136.

Bell, K.G., & Alvord, D.C. (1976). Pre-Silurian stratigraphy of northeastern Massachusetts, In: Page, L.R. (Ed.) contributions to the stratigraphy of New England. Memoir: Geological Society of America, 148, 179-216.

Bellon, H. (1988). Reconnaissance chronologique des deux premieres phases d‘activite volcanique en Dominique (Petites Antilles). CR Academy of Science Paris, 306, 1487-1492.

Billings, M.P. (1929). Structural geology of the eastern part of the Boston Basin. American Journal of Science, 5(18), 99-137.

Billings, M.P., & Tierney, F.L. (1964). Geology of the City Tunnel Extension. Journal of the Boston Society of Civil Engineers, 51, 111-154.

Bingen, B., Griffin, W.L., Torsvik, T.H., & Saeed, A. (2005). Timing of Late Neoproterozoic glaciation on Baltica constrained by detrital zircon geochronology in the Hedmark Group, south-east Norway. Terra Nova, 17(3), 250–258.

Blanc, F. (1983). Corrélations chronologiques et géochimiques des formations volcaniques du sud de la Basse Terre de Guadeloupe (Petites Antilles), University thesis, 171 pp., University of Grenoble, France.

Boudon, G., Semet, M.P., & Vincent, P.M. (1984). Flank failure-directed blast eruption at Soufriére, Guadeloupe, French West Indies: A 3,000-yr-old Mt St Helens? Geology, 12, 350- 353.

Bounden, G., Semet, M.P., & Vincent, P.M. (1987). Magma and hydrothermally driven sector collapses: the 3100 and 1, 500 Y.B.P. eruptions of the la Grande Découverte (la Soufriére)

245 volcano, Guadeloupe, French West Indies. Journal of Volcanological Geothermal Research, 33, 317-323.

Boudon, G., Le Friant, A., Komorowski, J-C., Deplus, C., & Semet, M.P. (1992). Volcano flank instability in the Lesser Antilles Arc: diversity of scale, processes, and temporal recurrence. Journal of Geophysical Research, 112, 28 pp.

Bouysse, P., & Guennoc, P. (1983). Données sur la structure de l‘arc insulaire des Petites Antilles, entre Ste. Lucie et Anguilla. Marine Geology, 53, 131-166.

Bouysse, P. (1984). The Lesser Antilles island arc: Structure and geodynamic evolution, In: Initial Reports of the Deep Sea Drilling Project, 78, 83-103. Washington, D.C., U.S. Government.

Bouysse, P., Andreieff, P., Richards, M., Baubron, J-C., Mascle, A., Maury, R.C, & Westercamp, D. (1985). Aves swell and northern Lesser Antilles ridge: rock dredging results from Arcante 3 cruise. In : A. Mascle (Ed.), Géodynamique des Caraïbes. Proc Symp Géodynamique des Caraïbes, Paris. Technip, Paris, 65–76.

Bouysse, P. (1988). Opening of the Grenada back-arc basin and evolution of the Caribbean Plate during the Mesozoic and early Paleogene, Tectonophysics, 149, 121-143.

Bouysse, P., Westercamp, D., & Andreieff, P. (1990). The Lesser Antilles Island Arc. Proceedings of Ocean Drilling Program, Science Results, 110, 29-44.

Bouysse P., & Mascle A. (1994) Sedimentary basins and petroleum plays around the French Antilles. In: A. Mascle (Ed.) Hydrocarbon and petroleum geology of France. European Association of Petroleum Geoscientists Special Publications, 4, 431–443.

Boynton, C. H., Westbrook, G. K., Bott, M.H.P., & Long, R. E. (1979). A seismic refraction investigation of crustal structure beneath the Lesser Antilles arc. Journal of Geophysical Society, 58, 371-393.

Brewer, T.S., Ahall, K-L., Menuge, J.F., Storey, C.D. & Parrish, R.R. (2004). Mesoproterozoic bimodal volcanism in SW Norway, evidence for recurring pre-Sveconorwegian continental margin tectonism, Precambrian Research, 134, 249-273.

Briden, J. C., Rex, D. C. Faller, D. M., & Tomblin J. F. (1979). K-Ar geochronology and palaeomagnetism of volcanic rocks in the Lesser Antilles. Philosophical Transactions of the Royal Society of London, A291, 485-528.

246

Bruckner, W.D. & Anderson, M.M. (1971). Late Precambrian glacial deposits in southeastern Newfoundland-a preliminary note. Geological Association of Canada Proceedings, 24, 95-102.

Bruckner, D. (1977). Significance of new tillites finds for east-west correlation of Proterozoic Avalon-zone formations in southeastern Newfoundland. Geology, 33, 357-364. Caldwell, D.W. (1964). The Squantum Formation: Paleozoic tillite or tilloid? In: J.W. Skehan (Ed.) Guidebook to Field Trips in the Boston Area and Vicinity. New England Intercollegiate Geological Conference, 56th Annual Meeting, 53-60.

Cameron, B., & Jeanne, R.A. (1976). New evidence for glaciation during deposition of the Boston Bay Group of southeastern New England. In: B. Cameron (Ed.). Geology of southwestern New England, New England Intercollegiate Geological Conference, 68th Annual Meeting, 117-134.

Carey, S.N., & Sigurdsson, H. (1980). The Roseau Ash: deep-sea tephra deposits from a major eruption on Dominica, Lesser Antilles Arc. Journal of Volcanology Geothermal Research, 7, 67-86.

Carey, S., & Sigurdsson, H. (1984). A model of volcanogenic sedimentation in marginal basins. Geological Society of London, Special Publications, 16, 37-58.

Carey, S., & Sigurdsson, H. (1987). Temporal variations in column height and magma discharge rate during the 79 A.D. eruption of Vesuvius. Bulletin: Geological Society of America, 99, 303-314.

Carey, S.N., & Sigurdsson, H. (1989). The intensity of Pinion eruptions. Bulletin of Volcanology, 51, 28-40.

Carlut, J., Quidelleur, X., Courtillot, V., & Boundon, G. (2000). Paleomagnetic directions and K/Ar dating of 0 to 1 Ma old lava flows from the La Guadeloupe Island (French West Indies): implications for time averaged field models. Journal of Geophysical Research, 105 (B1), 835- 849.

Carto, S. L. & Eyles, N. (2011) a. The Squantum Member of the Boston Basin, Massachusetts, USA. In: Arnaud, E., Halverson, G. P. & Shields-Zhou, G.A. (Eds.), The Geological Record of Neoproterozoic glaciations. Geological Society of London, Memoir, (in press).

Carto, S.L., & Eyles, N. (2011) b. The Gaskiers Formation, eastern Newfoundland, Canada. In: Arnaud, E., Halverson, G. P. & Shields-Zhou, G.A. (Eds.), The Geological Record of Neoproterozoic, Geological Society of London, Memoir, (in press).

Cas, R.A.F. & Wright, J.V. (1987). Volcanic successions, modern and ancient. Chapman and Hall, 1-528.

247

Cole, R.B., & Decelles, P.G. (1991). Subaerial to submarine transitions in early Miocene pyroclastic flow deposits, southern San Joaquin basin, California. Bulletin: Geological Society of America, 103(2), 221-235.

Condon, D.J., Prave, A.R., & Benn, D.I. (2002). Neoproterozoic glacial-rainout intervals: Observations and implications. Geology 30(1), 35–38.

Corsetti, F.A, Olcott, A.N., & Bakermans, C. (2006). The biotic response to Neoproterozoic snowball Earth. Palaeogeography, Palaeoclimatology, Palaeoecology, 232, 114-130. Crandel, D.R., Miller, C.D., Glicken, H., Christiansen, R.L., & Newhall, C.G. (1984). Catastrophic debris avalanche from ancestral Mount Shasta volcano, California. Geology, 12, 143-146.

Crowell, J.C. (1957). Origin of pebbly mudstones. Bulletin: Geological Society of America, 68, 993-1010.

Deplus, C., LeFriant, A., Boudon, G., Komorowski, J.C., Villemant, B., Harfod, C., Segoufin, J. & Cheminee, J.L. (2001). Submarine evidence for large-scale debris avalanches in the Lesser Antilles arc. Earth and Planetary Sciences Letters, 192, 145–157.

Dott, R.H. Jr. (1961). Squantum ―tillite‖, Massachusetts—evidence of glaciation or subaqueous mass movements? Bulletin: Geological Society of America, 72, 1289-1305.

Emerson, B.K. (1917). Geology of Massachusetts and Rhode Island. US Geological Survey, Bulletin, 597, 289 pp.

Evans, A.D. (2000). Stratigraphic, geochronological, and palaeomagnetic constraints upon the Neoproterozoic climatic paradox. American Journal of Science, 300, 347–433.

Eyles, N., & Eyles, C.H. (1989). Glacially-influenced deep-marine sedimentation of the Late Precambrian Gaskiers Formation, Newfoundland, Canada. Sedimentology, 36, 601-620.

Eyles, N. (1990). Marine debris flows: Late Precambrian ‗‗tillites‘‘ of the Avalonian– Cadomian orogenic belt. Palaeogeography, Palaeoclimatology, Palaeoecology, 79, 73–98.

Eyles, N. (1993). Earth‘s glacial record and its tectonic setting. Earth-Science Reviews, 35, 1 – 248.

Eyles, N., & Januszczak, N. (2004). ‗Zipper-rift‘: A tectonic model for Neoproterozoic glaciations during the break-up of Rodinia after 750 Ma. Earth Science Reviews, 65, 1-73.

248

Eyles, N. (2008). Glacioepochs and the supercontinent cycle after ~3.0 Ga: tectonic boundary conditions for global cooling. Palaeogeography, Palaeoecology, Palaeoclimatology, 258, 89- 129.

Fisher, R.V. (1971). Features of coarse-grained, high-concentration fluids and their deposits. Journal of sedimentology and Petrology, 41, 916-927.

Fisher, R.S., & Schmincke, H-U. (1984). Pyroclastic Rocks. Springer, Berlin Heidelberg New York Tokyo, 1-472.

Fisher, R.V. (1990) Transport and deposition of a pyroclastic surge across an area of high relief: The 18 May 1980 eruption of Mount St. Helens, Washington. Geological Society of America, Bulletin, 102, 1038-1054. Fisher, R. V., & Smith, G. A. (1991). Sedimentation in Volcanic Settings. SEPM Special Publication, 45, 257 pp.

Freundt, A. (2003) Entrance of pyroclastic flow into the sea: experimental observation, Bulletin of Volcanology, 65, 144-164.

Fujioka, K., Nishimura, A., Koyama, M., and Kotake, N. (1990). Bimodal arc volcanism and backarc rifting along Izu-Bonin Arc. Earth and Ocean Sciences, 71, 956.

Garcia, M.O. (1996). Turbidites form the slope failure on Hawaiian volcanoes, In: W.J. McGuire, A.P. Jones, & J. Neuberg (Eds.) Volcano Instability on the Earth and Other Planets. Geological Society Special Publications, 110, 281-294.

Glicken, H. (1986). Rockslide-debris avalanche of May 18, 1980, Mount St. Helens Volcano, Washington. Ph.D. dissertation, University of California, Santa Barbara, California, 303 pp.

Glicken, H., Komorowski, J-C., & Sheridan, M. (1989). Particle-particle interaction in the Mount St. Helens debris avalanche: Scanning Electron microscope Approach, Earth and Ocean Science, 70, p. 1423.

Glicken, H. (1991). Sedimentary architecture of large volcanic-debris avalanches. SEPM Special Publication, 45, 99-106.

Guest, J. E., Chester, D. K., & Duncan, A. M. (1984). The Valle del Bove, Mount Etna: its origin and relation to the stratigraphy and structure of the volcano. Journal of Volcanology and Geothermal Research, 21, 1-23.

Halverson, G., Hoffman, P.F., Schrag, D.P., Maloof, A.C., & Rice, A.H.N. (2005). Toward a Neoproterozoic composite carbon isotope record. Bulletin: Geological Society of America, 117, 1181-1207.

Hambrey, M.J., & Harland, W.B. (1981). Earth‘s Pre- Pleistocene Glacial Record. Cambridge:

249

Cambridge University Press, 1004 pp.

Hiscott, R.N. (1994).Traction-carpet stratification in turbidites- fact or fiction? Journal of Sedimentary Research, A64, 204-208.

Hoffman, P.F., Kaufman, A.J., Halverson, G.P., & Schrag, D.P. (1998). A Neoproterozoic Snowball Earth. Science, 281, 1342–1346.

Hoffman, P. F., & Schrag D. P. (2002). The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14(3), 129-155.

Hoffman, P.F. (2005). On Cryogenian (Neoproterozoic) ice-sheet dynamics and the limitations of the glacial sedimentary record. South African Journal of Geology, 108, 557–577.

Hoffman, P.F., & Li, Z.X. (2009). A palaeogeographic context for Neoproterozoic glaciation Palaeogeography, Palaeoclimatology, Palaeoecology, 277, 158–172.

James, K.H. (2006). Arguments for and against the Pacific origin of the Caribbean plate: discussion, finding for an inter-American origin. Geologica Acta, 4, 279-302.

Karig, D. E., and Moore, G. F. (1975). Tectonically controlled sedimentation in marginal basins. Earth and Planetary Science Letters, 26, 233-238.

Kirschvink, J.L. (2002). When All of the Oceans Were Frozen. Recherche, 355, 26–30.

Knoll, A. H., Walter, M.R., Narbonne, G.M. & Christie-Blink, N. (2006). The Ediacaran Period: a new addition to the geologic time scale. Lethaia, 39, 13-30.

Kokelaar, P. and Koniger, S. (2000). Marine emplacement of welded ignimbrite: the Ordovician Pitts Head Tuff, North Wales. Journal of the Geological Society, 157(3), 517 - 536.

Komorowski, J-C., Boudon, G., Semet, M., Villemant, B., & Hammouya, G. (2002). Recurrent flank collapses at Soufriére of Guadeloupe volcano: implications of acid hydrothermal fluids on edifice stability. Internal Congress, Montagne Pelee 1902-2002, Martinique island, May 12- 16, 69.

Komorowski, J-C., Boudon, G., Semet, M., Beauducel, F., Anténor-Habazac, Bazin, S., & Hammouya, G. (2005). In: M. Lindsay (Ed.) Guadeloupe, in Volcanic Atlas of the Lesser Antilles pp. 65-102, University of the West Indies, Seismic Research Unit, Trinidad.

Lacroix, J. (1904). La Montagne Pelee et Ses Eruptions. Masson, Paris, France.

250

Lavigne, F. (2000). Lahars: Deposits, origins, and behavior. Bulletin: Geological Society of France, 171(5), 545-557.

Leather, J., Allen, P., Brasier, M.D. & Cozzi, A. (2002). Neoproterozoic Snowball Earth under scrutiny: evidence from the Fiq glaciation of Oman. Geology, 30, 891– 894.

Lewis, J.F., & Robinson, E. (1976). A revised stratigraphy and geological history of the Lesser Antilles and the Mid-Atlantic Ridge. Deep-Sea Research, 16, 179-184.

Le Friant, A. (2001). Les destabilizations de flanc des volcans de l‘arc des Petites Antilles: origins et consequences. University thesis, 378 pp., University of Paris, Paris, France.

Le Friant, A., Bouden, D., Komorowski, J-C., Deplus, C. (2002). L‘île de la Dominique, á l‘origine des avalanches de debris les plus volumineuses del‘arc des Petites Antilles. CR Geoscience, 334, 235-243. Le Friant, A., Boudon, G., Komorowski, Heinrich, P.H. and Semet, M.P. (2003). Large-scale flank collapse events during the activity of Montagne Peleé, Martinique, Lesser Antilles. Journal of Geophysical Research, 108(B1), 2055.

Le Guen de Kerneizon, M., Carron, J.P., Bellon, H., & Maury, R.C. (1983). Saint Lucia (Lesser Antilles); distinction of principal magmatic series from petrochemical and geochronologic data. Bulletin: Geological Society of France, 25 (6), 845-853.

Lindsay, J.M., Stasiuk, M., & Shepard, J.B. (2003). Geological history and potential hazards of the late-Pleistocene to recent Plat Pays volcanic complex, Dominica, Lesser Antilles. Bulletin of Volcanology, 65, 201-220.

Lindsay, J.M. (2005). St. Lucia, in Volcanic Atlas of the Lesser Antilles, (Ed.) J.M. Lindsay, pp. 219-238, University of the West Indies, Seismic Research Unit, Trinidad.

Lipman, P.W., Normack, W.R., Moore, J.G., Wilson, J.B., & Gutmacher, E. (1988). The giant submarine Alika debris slide, Mauna Loa, Hawaii. Journal Geophysics Research, 93, 4279- 4299.

MacGabhann, B. A. (2005). Snowball Earth, the Dalradian Supergroup, and the Cryogenian- Ediacaran Boundary, 48th Irish Geological Research Meeting Programme and Abstracts 33.

251

Mascle, A., & Westercamp, D. (1983). Géologie d‘Antigua, Petites Antilles. Bulletin: Geological Society of France, XXV(6), 855–866.

Mattioli, G.S., Jansma, P.E., Jaramillo, L., & Smith, A.L. (1995). Sector collapse in island arc volcanoes: a digital topographic and bathymetric investigation of the Qualibou Depression, St. Lucia, Lesser Antilles, Caribbean. Journal of Science, 31, 163-173.

McCartney, W.D. (1967). Whitbourne Map Area, Newfoundland. Memoir: Geological Survey of Canada, 341, 1-133.

Meschede, M., Frisch, W. (1998). Non-Pacific origin of the Central American landbridge: Structural, geochronological, and paleomagnetic studies at the western and northwestern margin of the Caribbean plate. Trans. 14th Caribbean Geological Conference, Trinidad.

Miller, J.M.G. (1994). The Neoproterozoic Konnarock Formation, southwestern Virginia, USA: Glaciolacustrine facies in a continental rift, In: M. Deynoux, J.M.G. Miller, E.W. Domack, & N. Eyles (Eds.), Earth's Geologic Record; Cambridge University Press, Cambridge, p. 47-59.

Mimura, K., Kawachi, S., Fujimoto, U., Taneichi, M., Hyuga, T., Ichikawa, S., & Koizumi, M. (1982). Debris avalanche hills and their natural remnant magnetization, Nirasaki debris avalanche, central Japan. Journal of Geological Society of Japan, 88, 653-663.

Minster, J. B., & T. H. (1978). Jordan, Present-day plate motions. Journal of Geophysical Research, 83, 5331– 5354.

Monjaret, M.C. (1985). Contribution a l‘etude de l‘arc des Petites Antilles. Le Volcanisme de la Dominique. Donnes chronologiques, mineralogiques et geochimiques. Unpublished Thesis, Universite de Bretagne Occidentale, 1-77.

Mulder, T., & Alexander J. (2001). The physical character of subaqueous sedimentary density flows and their deposits. Sedimentology, 48, 269-299.

Murphy, J.B., & Nance, R.D. (1989). Model for the evolution of the Avalonian-Cadomian belt. Geology, 17, 735–738.

Murphy, J.B., van Staal, C.R., & Keppie, J.D. (1999). Middle to Late Paleozoic Acadian Orogeny in the northern Appalachians: A Laramide style plume-modified orogeny. Geology, 27, 653-656.

252

Narbonne, G.M., & Gehling, J.G. (2003). Life after snowball: the oldest complex Ediacaran fossils. Geology, 31, 27-30.

Narbonne G.M. (2004). Modular construction of early Ediacaran complex life forms. Science, 305, 1141–44.

Nichols, G. (1999). Sedimentology & Stratigraphy. Blackwell Science, Osney Mead, Oxford. 62-75.

O‘Brien, S.J., Wardle, R.J., & King, A.F. (1983). The Avalon Zone: a Pan-African terrane in the Appalachian Orogen of Canada. Journal of Geology, 18, 195-222.

O‘Brien, S.J., O‘Brien, B.H., Dunning, G.R., & Tucker, R.D. (1996). Late Proterozoic and related peri-Gondwanan rocks of the Newfoundland Appalachians, In: R.D. Nance, & R.D. Thompson (Eds.) Avalonian and related peri-Gondwanan terranes of the circum-North Atlantic: Boulder, Colorado. Geological Society of America Special Paper, 304, 9-28.

Olcott, A.N., Sessions, A.L., Corsetti, F.A., Kaufman, A.J., & de Oliveria, T.F. (2005). Biomarker evidence for photosynthesis during Neoproterozoic Glaciation, Science, 310, 471- 474.

Pazos, P.J., Sánchez Bettucci, L. & Loureiro, J. (2008). The Neoproterozoic glacial record in the Rio de la Plata Craton: a critical reappraisal. In: Pankhurst, M.J., Trouw, R.A.J., de Brito Neves, B.B. & de Wit, M.J. (Eds.) West Gondwana: Pre-Cenozoic Correlations Across the South Atlantic Region. Geological Society of London Special Publications, 294, 343–364

Perret, F. (1937). The Eruption of the Mt. Pelée 1929-1932. Publications, 458, 126 pp., Carnegie Institute, Washington.

Pe-Piper, G., & Murphy, J.B. (1989). Field trips in Nova Scotia. Atlantic Provinces Council on the Sciences, Pamphlet.

Picard, M., Schneider, J-C, & Boudon, G. (2006). Contrasting sedimentary processes along a convergent margin: the Lesser Antilles arc system. Geo-Marine Letters, 26 (6), 397-410.

Pindell, J. L., & Barrett, S. F. (1990). Geologic evolution of the Caribbean: A plate-tectonic perspective, In: G. Dengo, & J. E. Case (Eds.), The Caribbean Region: The Geology of

253

North America, Geological Society of America, Boulder, Colorado, p. 405-432

Pindell, J., Kennan, L., Stanek, K.P., Maresch, W.V., & Draper, G. (2006). Foundations of Gulf of Mexico and Caribbean evolution: eight controversies resolved. Geology Acta, 4, 303- 341.

Pinet, B., Lajat, D., Le Quellec, P., & Bouysse, P. (1985). Structure of Aves Ridge and Grenada Basin form multichannel seismic data. In: A. Mascle (Ed.) Géodynamique des Caraibes: Paris (Technip), 53-64.

Quinn, A.W., Ray, R.G., & Seymour, W.L. 1949. Bedrock geology of the Pawtucket Quadrangle Map GQ 1, scale 1:125,000.

Rankin, G. (1975). The continental margin of eastern North America in the southern Appalachians: the opening and closing of the Proto-Atlantic Ocean. American Journal of Science, 275, 298-336.

Rast, N. & Skehan, J.W. (1981). The geology of the Precambrian rocks of Newport and Middletown, Rhode Island, In: Hermes, O.D., and Boothroyd, J.C. (Eds.), Geologic field studies in Rhode Island and adjacent areas, 73rd New England Intercollegiate Field Conference Annual Meeting Guidebook, 67-92. Rast, N., & Skehan, J.W. (1983). The evolution of the Avalonian Plate. Tectonophysics, 100, 257-286.

Reid, R.P., Carey, S.N., Ross, & D.R. (1996). Late Quaternary sedimentation in the Lesser Antilles island arc. Bulletin: Geological Society of American, 108, 78-100.

Roobol, M.J., & Smith, A.L. (1980). Pumice Eruptions of the Lesser Antilles. Bulletin of Volcanology, 43(2), 277-286.

Roobol, M.J., Wright, J.V., & Smith, A.L. (1983). Calderas or gravity-slide structures in the Lesser Antilles Island Arc? Journal of Volcanology Geothermal Research, 19, 121-134.

Sayles, R.W., & LaForge, L. (1910). The glacial origin of the Roxbury Conglomerate. Science, 32, 723-724.

Sayles, R.W. (1914). The Squantum Tillite. Bulletin Harvard Museum Comparative Zoology, 66, 141-175.

Sayles, R.W. (1919). Seasonal deposition in the aqueo-glacial sediments. Memoirs of Museum Comparative Zoology, Harvard College, 47, 67 pp.

254

Schermerhorn, L. J. G. (1974). Late Precambrian mixtites: Glacial and/or non-glacial? American Journal of Science, 274, 673–824.

Schneider J-L. (2000) Volcaniclastic sedimentation in submarine settings: products and processes. In: H. Leyrit, & C. Motenat (Eds.), Volcaniclastic rocks, from magmas to sediments. Gordon & Breach, Amsterdam, pp 175–192.

Schneider, J-L., Le Ruyet, A., Chanier, F., Buret, C., Ferrière, J., Proust, J-N., & Rosseel, J-B. (2001). Primary or secondary distal volcaniclastic turbidites: how to make a distinction? An example from the Miocene of New Zealand (Mahia Peninsula, North Island). Sedimentary Geology, 145, 1–22.

Schrag, D.P., Erner, R.A., Hoffman, P.F., & Halverson, G.P. (2002). On the initiation of a snowball Earth. Geochemical Geophysics Ecosystems, 3, 1029-1036.

Schwab, F.L. (1981). Late Precambrian tallies of the Appalachians, In: M.J. Hambrey, & W.B. Harland, (Eds.), Earth‘s Pre-Pleistocene Glacial Record. Cambridge University Press, 751-755.

Shanmugam, G. (2000). 50 Years of the Turbidite paradigm (1950s-1990s): Deep-Water Processes and Facies Models-A Critical Perspective. Marine and Petroleum Geology, 17, 285- 342.

Shanmugam, G. (2002). Ten turbidite myths. Earth-Science Reviews, 58, 311–341.

Siebert, L. (1984). Large volcanic debris avalanches: characteristics of source areas, deposits, and associated eruptions. Journal of Volcanology of Geothermal Research, 22, 163-197. Siebert, L., Glicken, H. & Ui, T. (1987). Volcanic hazards from Bezymianny- and Bandai-type eruptions. Bulletin of Volcanology, 49, 435-459.

Sigurdsson, H. (1972). Partly-welded pyroclast flow deposits in Dominica, Lesser Antilles. Bulletin of Volcanology, 36, 148-163.

Sigurdsson, H., Sparks, R.S.J., Carey, S.N., & Huang, T.C. (1980).Volcanogenic sedimentation in the Lesser Antilles. Journal of Geology, 88, 523-540.

Sigurdsson, H., & Carey, S.N. (1991). Caribbean volcanoes: a field guide to Martinique, Dominica and St. Vincent. Geological Association of Canada, Toronto 1991 field trip guidebooks.

Skehan, J.W. & Rast, N. 1983. Relationship between Precambrian and other lower Paleozoic rocks of southeastern New England and other north Atlantic Avalonian terrains, In: Schenk, P.E. (Ed.) Regional trends in the geology of the Appalachian-Caledonian-Hercynian- Martitanide orogen. D. Reidel Publishing Company, 131-162.

255

Smith, A.L., & Roobol, M.J. (1980). Andesitic pyroclastic flows. In: R.S. Thorpe (Ed.), Orogenic andesites and related rocks. John Wiley and Sons, New York, N.Y.

Smith, G.W., & Socci, A.D. (1990). Late Precambrian sedimentary geology of the Boston Basin, In: A.A., Socci, S.J., Skehan, & G.W. Smith (Eds.), Geology of the composite Avalon terrane of southern New England. Geological Society of America Special paper 245, 75-84

Sohn, Y.K., Kim, S.B., Hwang, I.G., Bahk, J.J., Choe, M.Y., & Chough, S.K. (1997). Characteristics and depositional processes of large-scale gravelly Gilbert-type foresets in the Miocene Doumsan fan delta, Pohang basin, SE Korea. Journal of Sedimentary Research, 67, 130-141.

Sparks, R.S.J., Sigurdsson, H., & Carey, S.N. (1980). The entrance of pyroclastic flows into the sea, 1. Oceanographic and geological evidence from Dominica, Lesser Antilles. Journal of Volcanology Geothermal Research, 7, 87-96.

Speed, R.C., Smith-Horowitz, P.L., Perch-Nielsen, K.S., Saunders, J.B., & Sanfilippo, A.B. (1993). Southern Lesser Antilles Arc Platform: pre-Late Miocene stratigraphy, structure, and tectonic evolution. Geological Society of America Special Papers, 277, 1–16.

Tomblin, J.F. (1964). The volcanic history and petrology of the Soufriére Region, St. Lucia, Unpublished Ph.D thesis, Oxford University, U.K., 213 p.

Tomblin, J.F. (1965). The geology of the Soufriére volcanic centre, St. Lucia. Proceedings IV Caribbean Geological conference, Trinidad, 367-376. Tomblin, J.F. (1975). The Lesser Antilles and Aves Ridge, In: A.E.M. Nairn, & F.V. Stehli, (Eds.) The Ocean Basins and Margins, 3, New York, Plenum.

Ui, T., Suzuki, K., & Walker, G.P.L. (1983). Flow lineations of Koya low aspect-ratio ignimbrite, south Kyushu, Japan. EOS, 64, 876.

Urgeles, R., Masson, D.G., Canals, M., Watts, A.B., & Le Bas, T. (1999). Recurrent large- scale landsliding on the west flank of La Palma, Canary Islands, Journal of Geophysical Research, 104, 25331-25348.

Vincent, P.M., Bourdier, J.L., & Boudon, G. (1989). The primitive volcano of Mount Peleé: its construction and partial destruction by flank collapse. Journal of Volcanology Geothermal Research, 38, 1-15.

Voight, B., Glicken, H., Janda, R.J., & Douglas, P.M. (1981). Catastrophic rockslide avalanche of May 18, In: P.W. Lipman, & D.R. Mullineaux (Eds.) The 1980 Eruptions of Mount St. Helens. U.S. Geological Survey Professional Paper, 1250, 347-377.

256

Voight, B., Janda, R.J., Glicken, H., & Douglass, P.M. (1983). Nature and mechanics of the Mount St. Helens rockslide-avalanche of 18 May 1980. Geotechnique, 33, 243–273.

Webster, M.J. (1986). The structure of the Precambrian Newport Neck. The Lower Cambrian Pirate cove, and the East Passage formations, southeastern Rhode Island (M.Sc. thesis). Chestnut Hill, Massachusetts, Boston College, 213p.

Wadge, G. (1984). Morne Patates volcano, southern Dominica, Lesser Antilles. Geology Magazine, 122, 253-260.

Watts, A.B., & Masson, D.G. (1995). A giant landslide on the north flank of Tenerife, Canary islands. Journal of Geophysics Research, 100, 24487-24498.

Westercamp, D., & Tazieff, H. (1980). Martinique-Guadeloupe-Saint-Martin-La Désirade. In: Guides Géologiques Régionaux. Masson Paris, 135.

Westercamp, D., & Traineau. H. (1983). The Past 5,000 years of volcanic activity at Martinique (F.W.I): Implications for assessment of volcanic hazards. Journal of Volcanology and Geothermal Research, 17 159-185.

Westercamp, D., Pelletier, B., Thibaut, P.M., Traineau, H., & Andreieff, P. (1990). Carte géologique de la France (1/50 000), feuille Martinique. Orléans, BRGM.

Williams, H., & King, A.F. (1975). Southern Avalon, Newfoundland. Trespassey map area (IK). Papers Geological survey of Canada, 75(1), 11-15.

Williams, H., & King, A.F. (1979). Trepassey Map Area, Newfoundland. Geological Survey of Canada, Memoir, 389, 1-24.

Winterer, E. (1963). Late Precambrian pebbly mudstone in Normandy, France. In: A.E.M. Nairn (Ed.) Problems in Palaeoclimatology, 159-178. Interscience, London

Wohletz, K., Heiken, G., Ander, M., Goff, F., Vuataz, F.D., & Wadge, G. (1986). The Qualibou Caldera, St. Lucia, West Indies. Journal of volcanology Geotherm Research, 27 (1- 2), 77-115.

Wright, J.D., Boon, J.D., Green, M.O., & List, J.H. (1986). Response of the mid shoreface of the southern mid-Atlantic Bight to a ‗‗northeaster‘‘. Geo-Marine Letters, 6, 153–160.

Zartman, R.E. & Naylor, R.S. (1984). Structural implications of some radiometric ages of igneous rocks in southeastern New England. Bulletin: Geological Society of America, 95, 522- 539.

Zhao Y.Y., Zheng Y.F., & Chen F.K. (2009) Trace element and strontium isotope constraints on

257 sedimentary environment of Ediacaran carbonates in southern Anhui, South China. Chemical Geology, 265, 345-362.

Zhao Y. Y., & Zheng Y.F. (2010) Stable isotope evidence for involvement of deglacial meltwater in Ediacaran carbonates in South China. Chemical Geology, 271, 86-100.

258

CHAPTER 6:

SUMMARY

6.1 SIGNIFICANCE AND CONCLUSIONS

The objective of this thesis was to provide a direct sedimentological test of the

Snowball Earth hypothesis with regard to the Neoproterozoic Squantum Member of the Boston

Basin and other arc-related diamictite-bearing successions of the Avalonian-Cadomian

Terranes. It has shown that catastrophic global cooling and glaciation does not provide a useful explanation for environmental change and diamictite deposition along the Avalonian-

Cadomian Terranes at this time.

The principal findings of this thesis can be briefly stated as follows:

Field investigations focused on sedimentary facies analyses of outcrops of Squantum

Member show very clearly that the Neoproterozoic marine rocks of the Boston Basin consists of deep marine debrites (conglomerates and diamictites) and turbidites (sandstones and argillites) deposited by sediment gravity flows in a tectonically active, volcanically influenced arc basin within the Avalon Terrane. Continued assumptions that the Boston Basin rocks record the Gaskiers glaciation simply on the basis of the till-like character of diamictites is not supported by field data. The Squantum Member consists of debrites produced by mass flow where gravel and mud were mixed by downslope slumping; partially-mixed ‗heterogeneous‘ diamictite facies provide key evidence of this mixing process. Similarly, ‗ice-rafted horizons‘

(laminated pebbly argillites) are ‗co-genetic debrite-turbidite couplets‘ within low-density

259 turbidites that can be explained without any glacial or cold climate influence. This finding confirms the results of previous sedimentological work by Bailey (1984), which has been ignored by subsequent climate modelers.

Furthermore, field investigations from the Lesser Antilles Arc and Grenada Basin revealed that the depositional setting and principal facies of the Boston Basin as well as the other diamictite-bearing successions of the Avalonian-Cadomian Terranes can be explained in terms of modern volcanic depositional analogues without the need to invoke ‗catastrophic‘ environmental changes (e.g., ‗Snowball Earth Hypothesis‘; Hoffman et al., 1998; Evans, 2000;

Hoffman & Schrag, 2000, 2002; Kirschvink, 2002; Schrag et al., 2002; Rice et al., 2003;

Hoffman, 2005). The marine successions preserved within the basins of these ancient terranes closely resemble that of the Grenada Basin and Lesser Antilles Arc in the Caribbean. The dominant influence on sedimentation is volcanism and the release and reworking of very large volumes of volcaniclastic sediment into a steep-sided deep-water basin subject to recurring mass flow in the form of debris flows and turbidity currents. There is no case to be made for the theory that glaciers played a significant or even minor role in sedimentation in the basins of the Avalonian-Cadomian Terranes.

The Gaskiers glaciation is represented by marine rocks in Newfoundland; a terrestrial

(on-land) glacial sedimentary record is missing (Williams and King, 1979; Eyles and Eyles,

1989). The data from the modern glaciated volcano of Mount Rainier indicates that such facies are readily reworked and destroyed in the high-energy volcanic settings.

A glacial influence cannot be clearly recognized in the Neoproterozoic Avalonian-

Cadomian strata, which has traditionally been correlated with the Gaskiers Formation

260

(Thompson and Bowring, 2000; Narbonne and Gehling, 2003; Bingen et al., 2005; Halverson et al., 2005; Kawai et al., 2008; Hoffman and Li, 2009). One possibility is that these other rocks are of a different age and thus not strictly correlative or another possibility is that glaciation was limited in extent. Had the Gaskiers glaciation been as extensive as ‗Snowball

Earth‘ advocates believe, it is presumed that a widespread glacial influence on marine sedimentation facies would have been preserved throughout the Avalonian-Cadomian Belt.

The rare presence of a glacial signal in these basins is entirely inconsistent with the notion of global or near-global glaciation at this time.

Narbonne and Gehling (2003) argued that the short-lived (c. 580 Ma) Gaskiers glaciation (< 2 million years long) was a Snowball Earth event and had a major effect on the marine biosphere triggering the Cambrian ‗explosion.‘ This model is rejected on a number of counts. First, a ‗Snowball Earth‘ model is not supported by facies data and reconstructions of depositional environments presented here (see above). Second, given the restricted regional extent of ice covers during the Gaskiers glaciation it is not considered probable that climate cooling had any marked effect on marine environments. Third, new work by paleobiologists shows that diversification of skeletal animals began much later at 540 Ma (close the boundary of the Lower Cambrian at 542 Ma) and was much more gradual than previously thought, being drawn out over some 20 million years (see Maloof et al., 2010). Rifting of Rodinia, major paleogeographic re-organizations and attendant geochemical and sea level changes are now seen as more viable ‗uniformitarian‘ explanations. It is likely too, that glaciation is also a response to supercontinent rifting at this time (Eyles, 2008).

The overarching conclusion of this thesis is that paleoclimate models and their underlying assumptions must be tested against the rock record. Such models are thus

261 fundamentally reliant on field investigations of Neoproterozoic diamictites rocks using conventional facies analysis methods aimed at examining the sedimentology of these enigmatic rocks in their sedimentary, stratigraphic and tectonic context. This message is not new (Dott,

1961; Eyles, 1993; Crowell, 1999) but requires repeated emphasis.

6.2 OPPORTUNITIES FOR FURTHER RESEARCH

A comprehensive global evaluation of other glacial deposits linked to the Gaskiers glaciation found in China, Australia, Scotland, Urals, Norway, Arabian-Shield, Brazil, and

Russia is now required to fully evaluate the paleogeographic and tectonic setting of this event using detailed facies analysis methods. Furthermore, in light of the many climate models that stress the need for kilometer-thick, drifting ice covers on the world‘s oceans to reach ‗Snowball

Earth‘ conditions (known as ‗dynamic ice‘ or ‗sea glaciers‘; e.g., Goodman and Pierrehumbert,

2003; Poulsen, 2003; Lewis et al., 2003; Pollard and Kasting, 2005; Fairchild and Kennedy,

2007) there is also a demonstrable need for a global re-examination of other Neoproterozoic facies described as ‗ice-rafting debris‘ layers in order to better constrain the extent of ice during the Neoproterozoic.

The scope of such work would also include investigation and identification of possible triggering mechanisms on cooling, in particular the role of tectonics and the generation of high topography during the break up of Rodinia at the end of the Neoproterozoic.

262

REFERENCES

Crowell, J.C. (1999). Pre-Mesozoic ice ages: their bearing on understanding the climate system. Geological Society of America Memoir, 192, 106 pp.