Volcanic Debris Avalanches: From Collapse to Hazard 3030574105, 9783030574109

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Advances in Volcanology

Matteo Roverato Anja Dufresne Jonathan Procter   Editors

Volcanic Debris Avalanches From Collapse to Hazard

Advances in Volcanology An Official Book Series of the International Association of Volcanology and Chemistry of the Earth’s Interior – IAVCEI, Barcelona, Spain Series Editor Karoly Nemeth, Institute of Natural Resources, Massey University, Palmerston North, New Zealand

More information about this series at http://www.springer.com/series/11157

Matteo Roverato • Anja Dufresne Jonathan Procter

•

Editors

Volcanic Debris Avalanches From Collapse to Hazard

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Editors Matteo Roverato Department of Earth Sciences University of Geneva Geneva, Switzerland

Anja Dufresne Department of Engineering Geology and Hydrogeology RWTH Aachen University Aachen, Nordrhein-Westfalen, Germany

School of Earth Science, Energy and Environment Yachay Tech University Urcuqui, Ecuador Jonathan Procter School of Agriculture and Environment INR Massey University Palmerston North, New Zealand

ISSN 2364-3277 ISSN 2364-3285 (electronic) Advances in Volcanology ISBN 978-3-030-57410-9 ISBN 978-3-030-57411-6 (eBook) https://doi.org/10.1007/978-3-030-57411-6 © Springer Nature Switzerland AG 2021 This work is subject to copyright. All rights are reserved by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, expressed or implied, with respect to the material contained herein or for any errors or omissions that may have been made. The publisher remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. This Springer imprint is published by the registered company Springer Nature Switzerland AG The registered company address is: Gewerbestrasse 11, 6330 Cham, Switzerland

Editors acknowledge the unquestionable efforts of all researchers that contributed to the understanding of volcanic lateral collapses and debris avalanches before and after the seminal event at Mount St. Helens in 1980.

Foreword

If one could travel in time prior to the 1980 eruption of Mount St. Helens, one would realize that the understanding of volcanic debris avalanches was just in its infancy. At the time, the enigmatic textures and morphologies of the deposits from these avalanches were still surmised the product of other sedimentary processes, and there were even suggestions that the conical shape of hummocks were mere monogenetic magmatic features. Debris avalanches are among the most appealing volcaniclastic deposits due to their chaotic, varicolored textural features, abrupt downflow textural changes, presence of exotic fragments, and amazingly high mobility. Their emplacement has more enduring impact than the initial triggering volcanic process, inducing effects such as the formation of transient lakes whose deposits record subsequent volcanic episodes and paleoclimatic changes. Volcanic systems produce sediments at a spectacular rate, surpassing most if not all other geologic environments. Over geologic time, however, the record of single events is quickly obscured by erosion and new volcanic products. Unsurprisingly, amphitheaters produced by debris avalanches may be obliterated in a few thousand years by a pristine edifice produced by younger effusions. In most cases, the only evidence of a past edifice collapse is the remnant deposits, which may preserve hummocky topography. Recurrent collapses are unexceptional, further complicating the record. Therefore, discriminating among various collapse episodes is incredibly challenging and interpretations are commonly controversial. This volume is a collection of eleven impressive technical papers presenting the state of the art in volcano instability and debris avalanche deposits; the manuscripts reflect on the numerous publications on the matter over the last four decades, ranging from case studies to global comparisons utilizing multiple databases. The book explores fundamental aspects of the description, interpretation, and reconstruction of events that cause failure of volcanoes and the emplacement of debris avalanche deposits. The book begins with an historical review of debris avalanche studies prior to the 1980 Mount St. Helens eruption and continues with an overview of the terminology used to describe these phenomena, their size and geomorphological characteristics, the factors controlling volcano instability, and the possible climatic influence of their failure. Special attention is given to the sedimentology of debris avalanche deposits as a key to interpreting their origin and mobility, topics which are usually controversial. Computer simulations vii

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Foreword

are here presented as a complementary tool to detailed geological mapping, being useful in assessing hazard scenarios of stratovolcanoes with records of frequent collapse. The book is nicely complemented with the current understanding of lateral collapse processes in volcanic island settings and debris avalanche emplacement in submarine environments, including an overview of volcaniclastic ringplain sedimentation in an active volcanic setting. This book is the accomplishment of three young, enthusiastic, and highly qualified scientists, whom I wish to congratulate: They have ably put this fascinating yet catastrophic natural process in the spotlight. Although poorly known before the 1980 Mount St. Helens eruption, debris avalanches are now recognized as common events in the lifespan of stratovolcanoes. The editors assembled an amazing pool of scientists, who have devoted most (if not all) of their academic careers to understanding this awesome volcanic process. Given our growing awareness of the hazard that volcanoes pose on their surroundings, this contribution constitutes a very useful tool to volcanic monitoring, inasmuch as it encompasses our current understanding on debris avalanches as well as the unknowns that should be the focus of future studies. Lucia Capra CGEO-UNAM Juriquilla, Queretaro, Mexico

Contents

Volcanic Debris Avalanches: Introduction and Book Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Matteo Roverato and Anja Dufresne

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A Historical Perspective on Lateral Collapse and Volcanic Debris Avalanches . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lee Siebert and Matteo Roverato

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Terminology and Strategy to Describe Large Volcanic Landslides and Debris Avalanches . . . . . . . . . . . . . . . . . . . . . . . . . . Benjamin Bernard, Shinji Takarada, S. Daniel Andrade, and Anja Dufresne Distribution and Geometric Parameters of Volcanic Debris Avalanche Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Anja Dufresne, Lee Siebert, and Benjamin Bernard Factors Contributing to Volcano Lateral Collapse . . . . . . . . . . . . . Matteo Roverato, Federico Di Traglia, Jonathan Procter, Engielle Paguican, and Anja Dufresne

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Climatic Influence on Volcanic Landslides . . . . . . . . . . . . . . . . . . . 121 Gioachino Roberti, Nicholas J. Roberts, and Catherine Lit Volcanic Debris Avalanche Transport and Emplacement Mechanisms. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 Engielle M. R. Paguican, Matteo Roverato, and Hidetsugu Yoshida Sedimentology of Volcanic Debris Avalanche Deposits . . . . . . . . . . 175 Anja Dufresne, Anke Zernack, Karine Bernard, Jean-Claude Thouret, and Matteo Roverato Volcanic Debris-Avalanche Deposits in the Context of Volcaniclastic Ring Plain Successions—A Case Study from Mt. Taranaki . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 Anke V. Zernack Volcanic-Island Lateral Collapses and Their Submarine Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 255 Sebastian F. L. Watt, Jens Karstens, and Christian Berndt

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Computer Simulation of a Volcanic Debris Avalanche from Mt. Taranaki, New Zealand . . . . . . . . . . . . . . . . . . . . . . . . . . 281 Jonathan N. Procter, Anke V. Zernack, and Shane J. Cronin Cyclic Growth and Destruction of Volcanoes . . . . . . . . . . . . . . . . . 311 Anke V. Zernack and Jonathan N. Procter Glossary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 357

Contents

About the Editors

Matteo Roverato Born in 1981 in Como, Italy, Matteo Roverato studied geology at the University of Pisa (Italy). He completed his Ph.D. in 2012 at Centro de Geociencias, UNAM (Queretaro, Mexico) with a topic on Volcanic Debris Avalanches at Colima (Mexico) and Taranaki (New Zealand) volcanoes. Thanks to a prestigious Brazilian program for young scientists, he obtained a research fellowship at USP (Brazil) from 2013 to 2016, on Paleoproterozoic volcanic successions in the Amazon Craton. Afterward, he moved to Ecuador where he spent 3 years as Assistant Professor in Mapping, Field Methods for Geologists and Physical Volcanology. Currently, he is contracted as Maître-Assistant at the University of Geneva (Switzerland). Anja Dufresne Born in 1976 in the middle of Germany, Anja Dufresne gained her degrees in geology in Germany, US-Massachusetts, and New Zealand. She completed her Ph.D. at Canterbury University in Christchurch, New Zealand, where her work on volcanic and non-volcanic landslides around the world began. Returning to Germany, she was post-doc, lecturer, and project leader at the University of Freiburg. Since 2016, she works at the RWTH-Aachen University, where she is now assistant professor at the Department of Engineering Geology and Hydrogeology.

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About the Editors

Jonathan Procter Growing up in the shadow of Ruapehu and Taranaki in New Zealand, Jonathan Procter trained at Massey University, New Zealand. Jonathan completed his Ph.D. at Massey University in volcanology specializing in volcanic mass flows and the application of numerical simulations to inform volcanic hazard management. Associate Professor Procter is currently director of the Volcanic Risk Solutions group and manages the Geoscience group at Massey University. He currently leads the NZ National Science Challenge; Resilience, volcano research program and is active in finding new solutions to the hazards posed by our volcanoes.

Volcanic Debris Avalanches: Introduction and Book Structure Matteo Roverato and Anja Dufresne

Abstract

Volcanic debris avalanches (VDA) are, on the one hand, stunning natural phenomena, but, on the other, can pose serious threats to people and infrastructure. This first chapter aims to introduce a collection of themed papers gathered in a book, each illustrating the advancements of a different aspect of VDA research. As a state-of-the-art collection, the 11 papers provide a powerful tool for the volcanological community to enhance our understanding of their history and global distribution, collapse initiation and cyclic occurrence, problems with terminology, transport processes and deposit characteristics, climate impacts, application of numerical tools, and the records of marine and ringplain settings. The year 2020, marks the 40th Anniversary of the Mount St. Helens eruption and lateral collapse (Fig. 1), and we believe that it is time to collate our current knowledge around volcanic

M. Roverato (&) Department of Earth Sciences, University of Geneva, Geneva, Switzerland e-mail: [email protected] School of Earth Science, Energy and Environment, Yachay Tech University, Urcuqui, Ecuador A. Dufresne Engineering Geology and Hydrogeology, RWTH Aachen University, Aachen, Germany

debris avalanches. Volcanic debris avalanches (VDAs) are awesome natural processes forming by lateral collapse of volcanic edifices—they characterize almost all stratovolcanoes in different geodynamic settings around the world (Siebert 1984; McGuire 1996; Voight and Elsworth 1997; Voight 2000; Shea and Van Wyk de Vries 2010; Van Wyk de Vries and Davies 2015). Although less common, other volcanic structures such as dome complexes, calderas, monogenetic fields and shield volcanoes are also affected by this phenomenon. Gravitational collapse events like these are prominent during the Quaternary, and although identification of older deposits can be hindered by burial or erosional removal, largevolume VDAs have been documented throughout geologic time as far back as the Precambrian in the cratons of Australia (Trofimovs et al. 2004) and Brazil (Roverato 2016). Volcanic collapse events can occur cyclically, just once or at different spatial and temporal scales. Generally, the events involve large volumes (on the order of km3) of material, characterized by a frequency of thousands to tens of thousands of years (McGuire 1996; Capra et al. 2002; Zernack et al. 2012). Special cases, such as Mount St. Augustine Volcano (Alaska, USA) show, although with relatively “small” volumes (1000 km2 with an estimated original failure volume of 20 km3 and an inferred deposit volume of about 25 km3. Hummocks were found to the distal end of the deposit near the south Java coast, with distal hummocks preferentially stranded against six intervening Tertiary hill outliers and consisting dominantly of bedded pyroclastic-fall deposits. A deposit with similarities to a lateral-blast deposit locally overlies proximal parts of the avalanche deposit and a post-collapse cinder cone was constructed below the headwall scarp.

2.5 Japan 2.5.1 Bandai One of the most noted volcanic debris-avalanche deposits world-wide prior to 1980 was that

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generated by the collapse of Bandai volcano in central Honshu during a phreatic eruption nearly a century earlier. Seismologist Sekiya and geologist Kikuchi of the Imperial University of Tokyo arrived at Bandai-san only a few days after the July 15, 1888 eruption and published their classic 1889 detailed account of eyewitness descriptions and observations of the deposits. The eruption began with 15–20 rapid explosions on the north flank of Ko-Bandai volcano, one of several overlapping stratovolcanoes of the Bandai volcanic complex. Eyewitnesses observed the last explosion to be projected almost horizontally northward, followed by growth of a 4-km-high vertical eruption column that produced ash fall to the Pacific coast. A pyroclastic density current swept down Biwasawa valley on the east side, but similar deposits were not widely found to the north. Failure of the summit and north flank of Ko-Bandai produced a 1.5 km3 debris avalanche that buried several villages and traveled 11 km to the north, leaving a 1.5  2 km crater breached in that direction (Fig. 6). Sekiya and Kikuchi (1889) recognized that the greater part of KoBandai was not explosively ejected, but “thrown down much after the manner of a land-slip,” a conclusion reinforced by their seeing the failure of a 300-m-high section of the collapse scar and the subsequent progressive disintegration of the flowing mass. The curious mound topography “like so many miniature Fujiyamas” (Fig. 6) was considered to originate from syn- and postemplacement disintegration leaving conical shapes by forming talus aprons around them. The prominent reported explosions, however, led them to characterize the collapse scar as an explosion crater, noting its resemblance to the Valle del Bove at Etna and La Palma caldera in the Canary Islands, as well as to the Hoei crater formed on the flank of Fuji volcano during the 1707 eruption, the latter being an example of explosively generated craters without lateral collapse. Subsequent investigators invoked variable mechanisms for the Bandai collapse. Koto (1916) collected avalanche debris within a few days of the eruption but considered the debris to be the result of powerful steam explosions, disagreeing

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L. Siebert and M. Roverato

Fig. 6 Debris-avalanche hummocks of the 1888 collapse of Ko-Bandai volcano in the foreground with steam rising from fissure within source area. O-Bandai volcano lies

behind peak on source area rim at right (Sekiya and Kikuchi 1889)

with Begeat’s (1900) interpretation as a landslide independent of the volcanic activity. Textbooks variably interpreted the renowned Bandai eruption. Cotton (1944) described the Bandai deposits as a water-saturated mudflow, citing Jaggar’s (1930) view that the mounds represented the emplacement level of finer-grained mudflows that drained away. Rittmann (1962) considered the deposit to be that of a nuée ardente and Macdonald (1972) a volcanic mudflow transformed from an avalanche when encountering valley streams. Williams (1941) and Williams and McBirney (1979) recognized the avalanche origin of the Bandai deposits, although referring to their source as an explosion caldera. Detailed studies of the deposits of the 1888 eruption by Nakamura (1978) noted the apparent coincidence of the avalanche with the last explosion directed to the north and concluded that although the distal portion had transformed to a mudflow, the bulk of the deposit had traveled in a dry condition and introduced the term “volcanic dry avalanche.” Nakamura noted that little or no juvenile magmatic material was found in either the avalanche or air-fall tephra deposits. Moriya (1980), in his study noting the relation between hilly avalanche deposits and their breached source areas, referred to this process as Bandaian eruptions. Later work noted earlier

collapses at the Bandai volcanic complex, including the largest known, the late-Pleistocene Okinajima VDAD on the southwest flank of the Bandai volcanic complex, which was associated with a Plinian magmatic eruption and followed by construction of O-Bandai volcano in the collapse scar (Yamamoto et al. 1999). Chiba and Kimura (2001) mapped more than 13 VDADs from Bandai volcano, most of which were smaller volume than the 1888 Urabandai VDAD and the Okinajima VDAD. Magmatic eruptions ceased at Bandai prior to eruption of the voluminous ca. 25 ka Aira-Tn tephra from the Aira caldera in Kyushu (Yamamoto et al. 1999), with Chiba and Kimura (2001) noting stratigraphic evidence for multiple magmatic collapse events at Bandai prior to eruption of the Aira-Tn tephra with subsequent collapse events associated with deposits of phreatic eruptions.

2.5.2 Yatsugatake The late-Pleistocene Nirasaki debris-avalanche deposit from the Yatsugatake volcanic chain in central Honshu is the largest known in Japan. It extends more than 50 km from andesitic Gogendake at the southern end of the chain to a wedgeshaped segment bounded by the Shiro-kawa and Kamanashi-gawa rivers above Nirasaki city and beyond. Mason and Foster (1956) considered the deposit to be a mudflow and noted varying

A Historical Perspective on Lateral Collapse and Volcanic …

previous interpretations. Yamasaki (1898) and Misawa (1924) also considered the deposit to be a mudflow. Misawa (1924) inferred the conical hills on the deposit to be erosional remnants from stream flow of the Kaminashi-gawa river over the deposit. Ogawa (1932) mapped the deposit as that of a piedmont glacier, describing terminal moraines and some drumlins. Mason and Foster (1956) calculated a volume of >9.5 km3 including distal material removed by stream erosion. They inferred that the mudflow, which may have been hot, could have originated from an earthquake or possibly Peléan eruption that produced hot avalanches, with mobility of the mudflow enhanced by steam, crater-lake waters, melting snow, or rainwater. The origin of small hills 5–20 m in height in the broad piedmont section of the deposit and larger hills in the narrow more distal wedge rising above the surface of the deposit was discussed in detail. The authors considered and rejected several hypotheses for formation of the hills, including stream erosion remnants, compressional swells, hills remaining from adjacent subsidence, or hills stranded by the momentum of the flow on bedrock, lava flows, or other topographic highs. They considered that most of the hills originated when under hydrostatic pressure of the mudflow on the volcano's slope, material of relatively low viscosity from the interior of the mudflow was extruded through fractures in the drying, hardened crust. Hashimoto et al. (1976) mapped the deposit as the Nirasaki pyroclastic flow but discussed the characteristics of mudflow hills that rose above its surface. Shortly after 1980, Mimura et al. (1982) investigated the natural remanent magnetism of the more than 100 debris-avalanche hills as high as 80 m and as wide as 500 m, previously referred to as mudflow hills. They reported that the core of the hills consisted of individual megablock material that had been rotated chiefly in the horizontal plane and rafted down slope on the relatively low temperature and dry debris avalanche.

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2.5.3 Unzen The 1792 collapse of the dacitic Mayu-yama lava dome at the Unzen volcanic complex was Japan's most severe volcanic disaster. Although phreatic explosions had often been considered to have triggered the Mayu-yama (Mae-yama) collapse, Katayama (1974) documented that no eruptions occurred at Mayu-yama. He noted the lengthy descriptions of an eruption that began on February 10, 1792 at the neighboring volcano of Fugen-dake and considered it inconceivable that residents would not have mentioned eruptive activity closer to them at Mayu-yama. On April 21, more than 300 felt earthquakes formed fissures up to 1 km long and caused extensive damage in the coastal town of Shimabara. Loud detonations sounding to residents like guns from a Dutch frigate came from Mayu-yama, and rockfall produced dust clouds that at times obscured the mountain from Shimabara. Fear of a landslide prompted most of the residents to flee to the north, abandoning their homes. On April 29, part of the lava dome slowly slid 200 m eastward. By the middle of May, however, seismicity had subsided, and residents returned to their homes. At 8 pm on May 21, two intense earthquakes occurred, and 0.34 km3 of Tengu-yama, the southern of the two lava domes forming Mayuyama volcano, failed to the east; an avalanche was produced that swept into the Ariake Sea, extending the shoreline by almost 1 km and forming the Tsukumo-shima (Ninety-nine islands) (Fig. 7). A tsunami with three major wave crests swept over the most populated portion of the town and devastated a 77 km stretch of the Shimabara Peninsula coastline, causing 15,030 fatalities there and in provinces across the Ariake Sea. The landslide origin of the Mayuyama deposits of 1792 was recognized by volcanologists in Japan prior to 1980 (e.g., Ota 1969; Furuya 1974; Katayama 1974). Hoshizumi et al. (1999) noted earlier large-volume VDAs from Nodake and Myoken-dake of the Unzen volcanic complex.

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Fig. 7 Contemporary map of 1792 Unzen catastrophe. Hummocky terrain from a debris avalanche that swept into the Ariake Sea is shown in brown, with thin vertical line connecting proximal deposit area to top of the barren source area scar on Tengu-yama lava dome of Mayuyama volcano with the twin forested Shichimenzan lava dome to right unaffected by collapse. The onshore runup of associated tsunami that swept 77 km of the Shimabara peninsula coastline is shown in yellow with red line marking coastal road. Eruptive plume rises from Fugendake volcano (top center) with associated flank lava flow at center right; there was no eruptive activity at the Mayuyama dome complex. Image used with permission from Tokiwa Museum of Historical Materials, Shimabara

2.6 Kamchatka and Kurils 2.6.1 Bezymianny When ash plumes subsided to allow the first postcollapse images of Mount St. Helens in 1980, the new volcano profile revealed a remarkable resemblance to Bezymianny volcano on the Kamchatka Peninsula. Subsequent studies revealed the extent to which the 1980 collapse of Mount St. Helens was virtually a carbon copy of the 1956 eruption of Bezymianny. Both eruptions were preceded by a period of crypto-dome emplacement and the onset of phreatic and/or Vulcanian eruptions lasting as long as five months at Bezymianny. Catastrophic collapse and emplacement of a major debris avalanche was accompanied by a powerful directed explosion that swept a broad arc opposite the new breached crater, blowing down trees in an aligned manner, followed by vertical Plinian explosions with pyroclastic-flow emplacement.

L. Siebert and M. Roverato

Fig. 8 Bezymianny volcano before and after 1956 collapse. a Bezymianny volcano in 1946 (photo by B. Piip). b Post-collapse photo in May 1957 showing new lava dome in breached depression left by 1956 collapse (photo by G. Gorshkov). Images from Girina (2013)

Lava domes were subsequently emplaced in the newly formed breached craters (Fig. 8). Gorshkov (1959, 1963) and Gorshkov and Bogoyavlenskaya (1965) detailed the development of the eruption and distinguished its deposits, although with interpretations and terminology later clarified by work on the 1980 Mount St. Helens eruption. The Bezymianny eruption and its principal deposits were considered by Gorshkov (1963) and Gorshkov and Bogoyavlenskaya (1965) to be explosively generated, characterizing them as a Bezymianny-type directed blast, with deposits of a “directed-blast agglomerate” (later shown to be that of a debris avalanche) and a “directed-blast sand,” the latter identical to deposits of the finer-grained pyroclastic-density current (often referred to as a lateral-blast deposit) at Mount St. Helens. Gorshkov’s view of the Bezymianny deposits as the product of explosive eruptions was reinforced internationally in volcanological textbooks. Macdonald (1972) considered the Bezymianny

A Historical Perspective on Lateral Collapse and Volcanic …

eruption to be a type example of Peléean eruptions, with the explosions destroying the whole top of the mountain and glowing avalanches descending its flanks. Bullard (1976) likewise considered the “agglomerate flow” to be an ash flow. Peter Francis, who later led groundbreaking investigations of many VDADs in South America, considered the 1956 Bezymianny eruption a modern type example of Plinian eruptions in his Francis (1976) textbook. Williams and McBirney (1979) classified Bezymianny as a phreatic or steam blast eruption that destroyed the summit of the volcano and was followed by glowing avalanches. The interpretation of the 1956 Bezymianny deposits as explosively generated persisted in part after 1980. Melekestsev and Braitseva (1988) contrasted the deeper-seated Bezymianny crater of explosive origin to more shallow gravitational collapses at other volcanoes in Kamchatka and the Kurils.

2.6.2 Shiveluch A similar magmatic collapse event to that at Bezymianny took place at Shiveluch (also spelled Sheveluch) volcano in Kamchatka in 1964 (Gorshkov and Dubik 1970), although with significant differences from what occurred at Bezymianny. An earthquake at 7:07 am on November 12 triggered collapse of the edifice producing a 1.5 km3 debris avalanche that traveled 16 km from the summit, covering an area of 98 km2 and accompanied by minor phreatic explosions. At 7:20 am eruption of incandescent material began. A Plinian eruption column reached heights of about 15 km and continued for about an hour with the emplacement of pyroclastic flows that covered an area of about 50 km2. The 1964 eruption was considered to be another instance of an explosively generated directed blast eruption by Gorshkov and Dubik (1970), although in contrast to at Bezymianny, only deposits of the “lateral-blast agglomerate” were identified. Belousov (1995) attributed the lack of a “directed-blast sand” deposit to the absence of magma high in the edifice at the time of failure, with depressurization of the hydrothermal system in the upper edifice not being sufficient to trigger a magmatic pyroclastic

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density current (lateral-blast) eruption. Largescale lateral collapse was not unprecedented at Shiveluch; Ponomareva et al. (1998) and Belousov et al. (1999) documented eight or more collapse events at Shiveluch during the Holocene. All of the eight events documented by Belousov et al. (1999) were comparable to the 1964 eruption in that failure occurred prior to magma reaching the upper edifice and directedblast deposits were not documented. The Bezymianny, Shiveluch, and Mount St. Helens eruptions were instructive in the understanding of collapse events at volcanoes. Voight et al. (1981) in their work on the 1980 Mount St. Helens deposit and Ui (1983) and Siebert (1984) in assessing global analogs, noted the similarity of accounts of the 1956 Bezymianny and 1964 “directed-blast agglomerate” deposits to that of the Mount St. Helens debris-avalanche deposit. Ryabinin and Rodionov (1966) earlier had calculated that the Bezymianny edifice could not have contained sufficient steam to produce its 1956 crater, and Adushkin et al. (1984) showed that the 1956 airwaves were inconsistent with a blast of that size. Post-1980 field studies of the 1956 Bezymianny (Bogoyavlenskaya et al. 1985; Belousov and Belousova 1998) and 1964 Shiveluch deposits (Bogoyavlenskaya et al. 1985; Belousov 1995; Ponomareva et al. 1998) detailed their rockslide-debris avalanche origin. Belousov et al. (1999, 2007) noted distinctions between magmatic collapse events contingent on the location of magma within the edifice at the time of failure, with the absence of a directed-blast deposit at Shiveluch due to the lack of magma high in the edifice at the time of collapse, in contrast to conditions at Bezymianny and Mount St. Helens.

2.6.3 Kharimkotan Gorshkov (1970) extended the interpretation of lateral-collapse events as explosively generated directed blasts in Kamchatka to large breached craters in Kuril Islands volcanoes such as Kharimkotan (Harimkotan). A large VDA during the 1933 eruption of Kharimkotan (also known as Severgin) that extended the shoreline by a kilometer and generated a tsunami (Fig. 9) was

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L. Siebert and M. Roverato

Fig. 9 Kharimkotan (Harimkotan) volcano showing volcanic debris-avalanche deposits (VDADs). White arrows show direction of avalanche movement. The debris avalanches entered the sea, with the 1933 avalanche extending the shoreline and producing a tsunami that caused two fatalities on a neighboring island

(Belousova and Belousov 1995). Other Holocene debris avalanches include one about 1100 year BP on the east side underlying the 1933 VDAD and another at about 2000 year BP on the northwestern side (Belousov and Belousova 1996). Modified from 2006 Google Earth image

interpreted by Miyatake (1934) as a mudflow. The collapse was followed by Plinian eruptions that produced pyroclastic flows, but Belousov et al. (2007) considered magma to have been relatively deep beneath the summit at the time of collapse as the VDAD contained little juvenile material and, as at Shiveluch volcano in 1964, a lateral-blast deposit comparable to that at Mount St. Helens was not found. Five or more large VDAs were documented at Kharimkotan during the Holocene, including one to the east about 1100 year BP and another to the northwest about 2000 year BP that was also followed by Plinian eruptions (Belousova and Belousov 1995; Belousov and Belousova 1996).

a major geological event when during a multiyear grant from the National Geographic Society he discovered in 1916 the awe-inspiring Valley of Ten Thousand Smokes ignimbrite deposit near Katmai volcano. During his Katmai investigations Griggs encountered another puzzling deposit on the flanks of Mageik volcano that he considered of almost comparable interest to the Katmai crater and the Valley of Ten Thousand Smokes (Griggs 1920, 1922). In 1917 Griggs observed a massive chaotic deposit in the upper reaches of Martin Creek on the south side of Mageik volcano consisting of jumbled, fragmented rock of dominantly volcanic origin, but also containing sandstone blocks and segments of soil and plant remains. Many rock boulders were more than 10 m in maximum size, and the surface of the deposit contained numerous conical mounds and small ponds (Fig. 10). Griggs initially considered the Mageik deposit to be of glacial origin, noting the many circular ponds similar to kettle ponds on glaciers, but evidence of its sudden emplacement and morphological similarities to non-volcanic landslide deposits and that at Bandai volcano in Japan led to his interpretation of the deposit as originating

2.7 Alaska and Cascade Range 2.7.1 Mageik Robert F. Griggs was a University of Ohio botanist working on studies of kelp off the coast of the Alaska Peninsula in 1913 when he noticed the effect of ash from the seminal 1912 eruption of Novarupta (Katmai) on Kodiak Island vegetation. He was thrust into the role of dealing with

A Historical Perspective on Lateral Collapse and Volcanic …

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Fig. 10 Conical hummock (right) of Mageik debrisavalanche deposit in 1917 with person standing on top for

scale and flowing stream not yet incised through the 1912 avalanche deposit (Griggs 1920)

from a landslide on the flank of Mageik volcano. He discussed the origin of the mounds in some detail, noting that they had not previously been adequately described and concluded that they were characteristic features of large landslide deposits. The presence of Katmai (Novarupta) ash on the surface of the deposit and the lack of decay of entrained vegetation led him to consider the landslide to be contemporaneous with the 1912 eruption. Griggs calculated the volume of the deposit to be about 0.9 km3, assuming a conservative average thickness of about 10 yards (*9 m). Hildreth et al. (2000) mapped three VDADs originating from Mageik, the largest of which was about 0.35 km3 in volume. They noted that part of Griggs’ 1912 deposit was that of an older avalanche and estimated a volume of 0.05–0.1 km3 for the 1912 debris-avalanche deposit.

accounts (Dall 1884; Davidson 1884) focused primarily on observations of eruption plumes and a tsunami that swept across Cook Inlet, but also noted a ship captain’s observations that “from the summit a great slide of the mountain over half a mile broad had taken place towards the rocky boat harbor on the north-northwestward.” A geologic map of Augustine Island (Detterman 1973) considered parts of VDADs on the east, south, and northeast coasts to be in-situ lava flows but recognized “volcanic rubble flows” and mudflow deposits elsewhere. Kienle and Forbes (1976) attributed deposits and tsunami generation of the 1883 eruption to mudflows and nuées ardentes. Kienle and Swanson (1980) attributed flank hazards at Augustine mostly to lahars and pyroclastic flows entering the sea, which they revised in a Kienle and Swanson (1983) hazard assessment to note the presence of VDADs. The brief reference of Griggs (1920) to a landslide deposit on Augustine was in a paper on Mageik volcano, and it wasn’t until after 1980 that the debris-avalanche origin of widespread deposits at Augustine was recognized. Siebert et al. (1989, 1995) studied the 1883 Burr Point deposit and the larger West Island VDAD on the NW side and noted bathymetric evidence for other debris avalanches on all sides of the island. Begét and Kienle (1992) and Waitt and Begét (2009) expanded work on Augustine to investigate the extensive debris-avalanche deposits that ringed Augustine volcano. At least a dozen large VDAs

2.7.2 Augustine Prior to his work on the Mageik landslide deposit as part of his Katmai investigations, Griggs had briefly stopped at Augustine Island in Cook Inlet in 1913 during the first year of his Alaska studies (Griggs 1920, 1922). He observed the hummocky terrain at Burr Point on the northeast side of the island and later considered it as an analogue for his interpretation of the Mageik landslide deposit and others in the Katmai area. The latest debris avalanche at uninhabited Augustine volcano took place in 1883. Contemporary

26

L. Siebert and M. Roverato

Fig. 11 Augustine volcano has collapsed a dozen times or more in the past 2500 years. Debris-avalanche deposits in this image extend out to sea on all sides and are variably covered by deposits of younger eruptions closer

to the volcano. Selected deposits are named, with the 1883 AD Burr Point deposit (in red) being the youngest. VDADs mapped by Waitt and Begét (2009) are overlaid on July 3, 2018 Landsat 8 imagery

were found to have occurred within the past 2500 years, making Augustine the volcano with the highest-known frequency of edifice-failure debris avalanches (Fig. 11).

noted an impressive array of previous interpretations for the deposit. Diller et al. (1915) had considered the hills of lava and tuff to originate at least in part from minor local eruptions that pierced Cretaceous bedrock, forming small volcanic deposits around individual vents. Fenner (1923) proposed that a shallow sill had been intruded beneath Shasta Valley, with small bodies of magma breaching the surface to form the hills. Williams (1949) and later Mack (1960) and Hotz (1977) mapped the proximal part of the deposit as glacial moraines originating from a glacier on the northwestern slopes of the volcano and flat-lying areas between hummocks as fluvioglacial outwash. Distal portions were described as Tertiary Western Cascade lavas and tuff breccias with intervening alluvium, with hillocks representing erosional remnants.

2.7.3 Shasta The hilly topography of the massive VDA from Shasta volcano in northern California, the largest Quaternary debris avalanche known in the western U.S., puzzled geologists for more than a half century. Its origin remained a mystery until after the 1980 eruption of Mount St. Helens, when Harry Glicken and other USGS volcanologists driving across the extensive deposit on Interstate5 highway to and from destinations in California from the Cascade Volcano Observatory noted its resemblance to what they had been working on at Mount St. Helens (Fig. 12). Crandell (1989)

Fig. 12 Hummocks of Shasta Valley debris-avalanche deposit with Shasta volcano in the background. Large feature on right horizon is not a hummock, but Holocene Black Butte lava dome. Photo Stephen Brantley

A Historical Perspective on Lateral Collapse and Volcanic …

Christiansen (1982) briefly noted the presence of a large debris-avalanche deposit in the Shasta Valley, and the deposit was studied in detail by Crandell et al. (1984) and Crandell (1989). Crandell (1989) mapped a deposit that covered at least 675 km2 with a volume of 45 km3 or more that was emplaced between about 300,000 and 380,000 years ago. The northern terminus of the hilly block facies lies about 49 km from the present summit of the volcano, with intervening flat areas of matrix facies, consisting of unsorted and unstratified mudflow-like deposits chiefly from the volcano. Shasta displays some of the largest hummocks of subaerial debris avalanches, with some reaching more than a kilometer in length. Analysis of hummock dimensions (Herrick et al. 2013) showed that Shasta hummocks displayed the same logarithmic decay in hummock size with distance from the source that was apparent in data sets with the smallest-diameter hummocks as longer travel distances provided more time for disaggregation of block-facies hummock material (Fig. 13).

2.7.4 Chaos Crags A much smaller debris-avalanche deposit (0.15 km3) at Chaos Crags lava-dome complex in the Lassen volcanic center of northern California was recognized by Williams (1928), Heath (1960) and Crandell et al. (1974) as a rockfallavalanche deposit. Williams (1928) considered the avalanche to have required a basal wet component that moved as a mudflow to explain its great mobility. Heath (1960) considered the three discrete lobes of the deposit to have been emplaced during separate events spanning as much as 1200 years, although Crandell et al. (1974) documented evidence for a single retrogressive failure that took place about 300 years ago and calculated a minimum velocity of 160 km/h based on runup of 120 m of the distal part of the avalanche on to the flanks of Table Mountain. Eppler et al. (1987) modeled avalanche kinematics with respect to potential hazards impacting a Lassen National Park road and visitor center constructed on the deposit and concluded that a future avalanche would be of

27

smaller size given the reduced volume of the dome and potential to reach the road but not visitor facilities.

2.8 Mexico Colima Volcano is considered the most active volcano in Mexico and is also known as Volcán de Fuego (Fire Volcano). The volcano was built on the southern flank of the older Nevado de Colima volcano, and inside the Paleofuego depression, a 4-km-wide relic of a Holocene lateral collapse (Robin et al. 1987; Luhr and Prestegaard 1988). This depression has been interpreted in different ways since the beginning of the last century. Waitz (1906) was the first to write about the origin of the scar of Colima volcano, considering it as a maar and defending his ideas over the next decades (Waitz 1932). Mooser (1961), on the other hand, defined the scar as a “submergence caldera” and hypothesized that the depression was generated by multiple phases of magma-chamber collapse. Mooser also described, for the first time, two other “calderas” of the neighboring Nevado de Colima volcano. The same hypothesis was supported by Luhr and Carmichael (1980), who defined these depressions as summit calderas likely formed by the collapse of shallow magma chambers. One year before, Demant (1979) considered the theories of Waitz (1906, 1932) as inconsistent with a maar formation and suggested, instead, that the depression of Colima volcano had formed through cyclical violent ash and pumice eruptions. In 1978 the Instituto Nacional de Estatística, Geografia y Informática (National Institute of Statistics and Geography) started a geological survey of Mexico, resulting in a series of geological maps of Colima volcano and surrounding areas (INEGI 1984). In these maps, the VDADs cropping out in the volcano ring-plain were considered as volcanic breccias. In the early 1980s, the “caldera” of Colima volcano was finally recognized to be a debris avalanche-related collapse scar (Luhr and Carmichael 1982; Robin et al. 1984). Several years

28

L. Siebert and M. Roverato

Fig. 13 Map of 487 larger-size hummocks of Shasta Valley VDAD with red dashed line showing inferred

source area (Herrick et al. 2013) and red asterisk marking location of Fig. 12 photo

later, Robin et al. (1987) and Luhr and Prestegaard (1988) described the collapse event in detail and the associated debris avalanche as a single, broad deposit although with divergent ages of 9370 ± 400 year BP and 4280 ± 110 year BP, respectively. A geomorphology study (Hubp et al. 1993) of the Colima area describes the hummocky topography in the SE and SW Colima volcano ring plain as

possibly related to two distinct VDADs, using the Robin/Luhr ‘s age divergence to support their idea. The Colima volcanic complex is now known to have a complex stratigraphy of collapse events, with several Pleistocene events at ancestral Nevado de Colima volcano (Cortes et al. 2010) and as many as eight major deposits from Paleofuego and the historically active Volcán de

A Historical Perspective on Lateral Collapse and Volcanic …

Fuego volcanoes during the past 30,000 years, the most recent of which was the 2500 year BP El Remate-Armería VDAD (Stoopes and Sheridan 1992; Komorowski et al. 1997; Capra et al. 2002; Cortes et al. 2005, 2010, 2019; Roverato et al. 2011; Roverato and Capra 2013).

2.9 South America 2.9.1 Chimborazo Massive 6263-m Chimborazo volcano, Ecuador’s highest peak, has undergone a major collapse event that remained unrecognized until the late 1980s. Deposits of the largest Chimborazo collapse underlie the city of Riobamba, part of a VDA (Riobamba Formation) that travelled more than 40 km to the southeast (Alcaraz 2002; Bernard et al. 2008) (Fig. 14). In the early 1980s the Riobamba VDAD was considered a lahar deposit caused by an eruption of Chimborazo inferred to be during or shortly after the last glaciation (Clapperton 1983). The remnant of the collapse scar is still slightly visible in the northwestern volcanic flank, although later volcanic activity and glacial erosion have obliterated most of it. Kilian (1987) and Kilian et al. (1995) described the scar as a subsidence caldera, although no associated ignimbrite deposits have been found, and defined the Riobamba formation as generic volcanic debris. Beginning in the late 1980s several authors recognized the edifice collapse and associated VDAD (Beate and Hall 1989; Clapperton 1990; Beate et al. 1990), although there have been differing views regarding which of the edifices of the Chimborazo volcanic complex sourced the avalanche and the age of collapse. Barba et al. (2005) and Samianego et al. (2012) used geochemical comparisons of edifice and deposit rocks and avalanche deposit size considerations to propose that the collapse originated from the basal CH-I edifice about 60–65 ka and that the three current summits post-date the avalanche with the previously hypothesized subsidence caldera scar originating from the debris-avalanche event. The first detailed studies of the Riobamba VDAD by Alcaraz (2002) and Bernard et al. (2008) showed

29

that the avalanche deposit covers an area of about 280 km2 with an estimated volume of >11 km3 (Bernard et al. 2008), making it one of the largest known in Ecuador. The Riobamba VDAD is dominated by block facies material, although the relatively small size of larger blocks (100

Inokuchi (1988)

Madeira Is

8

Quartau et al. (2018)

Japan-Hokkaido

13

Yamagishi (1996)

Canary Is

15

Krastel et al. (2001)

Japan

128

Inokuchi (2006)

Canary Is

11

Masson et al. (2002)

Japan

67

Yoshida (2016)

Canary Is

26

Acosta et al. (2004)

Kurile Is

>40

Belousova and Belousov (2011)

Canary Is

31

Hunt et al. (2014)

Kamchatka and Kurils

13

Melekestsev and Braitseva (1988)

Cape Verde Is

8

Masson et al. (2008)

Kamchatka

33

Ponomareva et al. (2006)

Global

43

Ui (1983)

Aleutian Is

14

Coombs et al. (2007)

Global

97

Siebert (1984)

Aleutian Is

17

Montanero and Beget (2011)

Global

195

Siebert et al. (1987)

Cascade Range

48

Siebert and Vallance (2017)

Global

56

Holcomb and Searle (1991)

Hawaiian Is

11

Moore et al. (1989)

20

Mitchell (2003)

Global (Rift islands) Hawaiian Is

20

McMurtry et al. (2004)

Global

316

Dufresne et al. (2008)

Hawaiian Ridge

>68

Moore et al. (1994)

Global

301

Bernard (2008)

Society and Austral Is

13

Clouard and Bonneville (2004)

Global

182

Blahůt et al. (2019)

Mexico

23

Capra et al. (2002)

Global

1001

This study; Dufresne et al. (2020a)

Submarine event totals exclude debris flows or low-velocity slumps. Blue color indicates inventories with largely submarine deposits. See Dufresne et al. (2020a—this volume) for chronologically sequenced list with additional context

VDADs have been identified as far back as the Precambrian (Trofimovs et al. 2004; Roverato 2016). Recent collapse events >0.1 km3 in volume have averaged more than 5 per century since 1500 AD and about 7 per century since 1800 AD (Table 2). Comparable events with volumes 437

T

M

3

Grilli et al. (2019)

Stratovolcano

2004

0.2

–

32

A

–

–

Tsuchiya et al. (2009)

Galapagos

Caldera

1988

0.9

10

–

–

M

2

Chadwick et al. (1991)

St. Helens

Cascades

Stratovolcano

1980

2.5

64

57

P,A,L

Mb

5

Voight et al. (1981)

Shiveluch

Kamchatka

Dome

1964

1.5

98

–

–

M

4

Gorshkov and Dubik (1970)

Volcano

Location

Type

Anak Krakatau

Indonesia

Stratovolcano

Bawakaraeng

Indonesia

Fernandina

Year

Bezymianny

Kamchatka

Stratovolcano

1956

0.8

30

–

–

Mb

5

Gorshkov (1959)

Kharimkotan

Kurile Is

Dome

1933

0.5

>20

2

T

M

5

Belousova and Belousov (1995)

Hakuba-Oike

Japan

Compound

1911

0.15

–

–

–

–

–

Yoshida (2016)

Ritter Island

Melanesia

Stratovolcano

1888

(2.44.2)

100

3000?

T

M

2?

Johnson (1987)

Bandai

Japan

Stratovolcano

1888

1.5

34

461

A,P

P

4

Sekiya and Kikuchi (1889)

Augustine

Alaska

Dome

1883

0.3

21

–

–

M

4

Siebert et al. (1995)

Krakatau

Indonesia

Stratovolcano

1883

(3.8)

–

(?)

T,P

M

5

Camus et al. (1992)

Sinarka

Kurile Is

Stratovolcano

1878

0.5?

–

?

A

M

4

Belousova and Belousov (2011)

Tate-yama

Japan

Stratovolcano

1858

0.2

–

Many

–

–

–

Nozaki (2015)

Suwanosejima

Japan

Stratovolcano

1813

>1.0

–

–

–

M

4

Shimano et al. (2013)

Tutupaca

Peru

Stratovolcano

1802?

0.7?

–

–

–

M

4?

Samaniego et al. (2015)

Unzen

Japan

Dome

1792

0.34

15

15,030

T,A

–

–

Ota (1969)

Asama

Japan

Stratovolcano

1783

0.14

n/d

1491?

A,P,L

M

4

Tamura and Hayakawa (1995)

Papandayan

Indonesia

Stratovolcano

1772

0.14

18

2957

A

P

3

Glicken et al. (1987)

OshimaOshima

Japan

Stratovolcano

1741

2.5

69

1475

T

M

4

Satake and Kate (2001)

Augustine

Alaska

Dome

1700?

0.15

10

–

–

M

?

Siebert et al. (1995)

Nabukelevu

Fiji

Dome

1650?

>0.1?

–

–

–

M

?

Cronin et al. (2004)

Callaqui

Chile

Stratovolcano

1630?

0.5

–

636

L

M

4

Thouret et al. (1990)

Dome

1540?

0.5

30

–

–

Mb

4?

Siebert et al. (1995)

Stratovolcano

1500?

0.23

–

–

–

–

–

Scott et al. (2001)

Volcano

Location

Type

Komagatake

Japan

Stratovolcano

Ruiz

Colombia

Stratovolcano

Augustine

Alaska

Rainier

Cascades

Year

Fatality agents: A, avalanche; L, lahar; P, pyroclastic density current; T, tsunami. Known tsunami fatalities at Krakatau in 1883 are not listed because, although large-volume submarine debris-avalanche deposits have been identified (Camus et al. 1992; Deplus et al. 1995), the number of potential related tsunami fatalities is not known. Type of collapse: M, magmatic eruptions; Mb, magmatic eruption with lateral blast; P, phreatic eruption. At Unzen volcano in 1792 there was no eruption at the Mayuyama dome complex where collapse occurred although there was at neighboring Fugen-dake volcano (Fig. 2.7). VEI: Volcanic Explosivity Index (Newhall and Self 1982). Single reference for each deposit focuses on early work pertaining to edifice collapse; additional references can be found in text of selected case studies

When magma is even deeper, phreatic eruptions where no juvenile magma reaches the surface can occur, as at Bandai in 1888. Alternatively, failure can occur in the absence of eruptive activity, as in about a fifth of the events tabulated here. This degree of magmatic involvement cannot be extrapolated to older events, however, as associated eruptive activity is documented for only a small fraction of earlier collapse events (Dufresne et al. 2020a—this volume). It should be noted that of collapses an order of magnitude smaller in volume than those documented in Table 2 during this same 500-year interval, more than half were non-eruptive events. Fatalities are known to have occurred at more than half of these events listed in Table 2 and would have occurred at many more had they not occurred in uninhabited or sparsely populated regions. The total number is strongly dominated by the debris-avalanche induced tsunamis during the Mayu-yama collapse at Unzen volcano in 1792. These numbers could be substantially higher depending on the uncertain degree of tsunami generation associated with an apparent debris-avalanche component of the 1883 Krakatau eruption (Camus et al. 1992; Deplus et al.

1995). This type of tsunami hazard was recently underscored by the collapse of Anak Krakatau in December 2018 during early stages of work on this volume (Grilli et al. 2019).

4

Discussion

4.1 Deposits The generation and emplacement of lateral edifice-collapse debris avalanches has been the last mechanism of large-scale edifice destruction and subsequent volcaniclastic processes to be widely recognized and understood. These deposits of often hilly volcanic material located far beyond the volcanoes themselves had long puzzled geologists. Deposits at the smaller end of the size spectrum of large VDAs (0.3 g) occurs (Deutschle 2013). Della Pasqua et al. (2016) undertook a similar analysis to understand the hazards posed by VDAs in the Taranaki region and found that susceptibility to failure will increase when there is higher seismic loads on weak, stratified layers. Alloway et al. (2005) identified a number of faults that traverse the Mt. Taranaki edifice and reconstructed earthquakes of magnitude 6.7–7.2 with shaking intensities of magnitude >8 with a recurrence interval of 4 ka. Keefer (1984) suggests that rock and soil avalanches require strong shaking intensities of magnitude >7–8. It is conceivable that a number of Mt. Taranaki VDADs were initiated by ground shaking associated with movements on those faults, but no conclusive evidence is presented in the literature. Earthquakes triggering VDAs is well recorded and while the triggering effects of the ground shaking to induce collapse have been elucidated from landslide investigations through understanding the physical properties of soils and rock, the initiation of the mass that suddenly loses its stability and move downslope is not well constrained or modelled accurately for volcanic edifices.

M. Roverato et al.

As we can see, most of the information available in the literature regarding eruptions and earthquakes, and their role in sector collapses, arises from historical cases or is inferred by theoretical models and simulations. However, if we take a look at the statistics of real cases, there is not much pre-historic information available about the triggering mechanisms that promote a sector collapse. Dufresne et al. (2020, this volume) reports that only 16% of VDADs in a new global database of 1001 events have a related triggering mechanism; most of them are associated with magmatic activities. It is obvious that it proves difficult to associate a volcanic collapse with a specific trigger. However, some information on factors and processes involved in past collapses is recorded in the resulting deposits. In Table 3, we summarize key features that could be useful for studying VDADs in the field by providing information on the state of the pre-collapse volcanic edifice and the processes that acted before or during a sector collapse. Efforts to couple lateral collapse models and flow models as proposed in Voight and Sousa (1994) need to be further developed. Being able to transfer those simulations and predictions into a hazard and risk analysis framework (e.g. Mead et al. 2018) to provide a more holistic analysis of the event is also required considering that much of the motivation for understanding volcanic debris avalanches arose from better understanding the hazardous events that occurred over the last 300 years.

5

Conclusions

Volcanic sector collapses can be prepared by a wide variety of destabilizing factors, which are followed by triggering elements that drive the collapse (Fig. 7). Although we group these factors in distinctive sections, it is evident that most of them are inherently related. In fact, factors that prepare and trigger volcanic sector collapses are not simple but can be multiple and cumulative. Over-steepened slopes, magma intrusions, hydrothermal activity, climate fluctuations, deformation of the basement, eruptions,

Factors Contributing to Volcano Lateral Collapse

111

Table 3 Key features in a VDAD that could suggest what happened in the volcano edifice before and/or during the lateral collapse event Key features

Pre/syn-collapse setting

Hydrothermal alteration

The VDAD displays: • color stains in those reaches (e.g. proximal) where the volcaniclastic material is not mixed with other materials • homogeneous yellowish matrix in portions (e.g. medial/distal) where the material is partially or completely mixed • high clay content

The edifice was affected by hot fluid circulation favoring hydrothermal alteration responsible for the weakening of the strength of the rock and promoting instability of the volcano

Water content

The VDAD is characterized by a watersaturated (or semi-saturated) aspect displaying facies with characteristics between debris-flow deposits (DFDs) and VDADs • homogenous and partially cemented matrixpreservation of millimetric porosity The VDAD is often associated to DFDs that lying directly on top of it

The edifice was affected by high water content due to different climatic conditions such as: • high precipitation in tropical settings • presence of snow and/or ice in boreal settings or at high altitudes in tropical settings These factors contribute to the instability of the volcanic edifice and could also increase the area affected by a VDA through flow transformation into more mobile debris flows

Pyroclastic deposits

Pyroclastic deposits such as pyroclastic density current (PDC) or fall-out deposits lying directly in contact on top of the VDAD displaying sharp contacts

Pyroclastic material situated on top of the VDAD suggests that an eruption likely occurred after the collapse event as consequence of the depressurization of the magmatic system. The sharp contact between PDC and VDA deposits suggests that the VDA had already ceased to move when the PDC material emplaced. The magmatic event was likely not responsible for the collapse and other triggering mechanisms are involved

PDC deposits such as directed blasts and/or blast-generated PDC deposits are “intimately” associated with the VDAD. Some characteristics are: • the contact between PDC and VDA deposits is very irregular • twisted clastic dikes of PDC material penetrate into the VDAD • in distal reached PDC deposits are located below and above the VDAD These features could be associates to the following section (juvenile material)

These characteristics indicate that the VDA was still moving when the PDC material started to emplace on the upper surfaces of the VDA. PDC deposits located below the VDAD in distal portions suggest that the kinetic energy of the PDC material was higher than the VDA and that it emplaced before the arrival of the collapsing mass. This section could be associates with the following one (juvenile material)

Fresh fragments of volcanic material (lava fragments or other pyroclastic material) are immersed into the VDAD

It indicates that the collapse was associated to a magmatic event that triggered the collapse. Fresh lava fragments could indicate the presence of a summit dome or a shallow intrusion. Vesiculated juveniles are the product of the explosive character of the eruption

Juvenile material

112

M. Roverato et al.

Fig. 7 Sketch representation of the instability factors that can destabilize a volcano and the triggering mechanisms, such as explosive activity and earthquakes, which promote a volcanic sector collapse. The top right rectangle represents schematically how the hydrothermal activity affects a volcanic edifice, promoting leaching and

dissolution of the rock, favoring the formation of voids and the migration and accumulation of clay minerals, and consequently, fluid storage and impermeability. This promotes loss of strength and shear reduction of the rock. See the text for more details

earthquakes, are just a list of elements that could promote a volcanic collapse. Although it is necessary to identify the differences between those factors that define the degree of instability from those mechanisms that trigger the collapse, it is also important to define those factors that, in combination, could lead to different catastrophic scenarios. Indeed, the mechanism that triggers a sector collapse could be independent of the causes that prepare its instability. But it is also true that multiple factors in consort could drastically enhance the power and magnitude of the collapse event and affect its runout. The intrusion of a magma body (e.g. cryptodome) into a volcanic edifice, for example, has multiple consequences for the stability of the volcano. On one hand, the magma body deforms the edifice structure and, if the magma reaches the surface, its products could overload the volcanic flanks (destabilizing factor). On the other hand, the hydrothermal fluids associated with the magma body could promote all those processes

of rock leaching and dissolution that weaken the volcanic material. If the intrusion of a magma body, linked with an earthquake or volcanic tremors, could be considered a factor that induce a volcanic collapse (trigger), clay mineral formation and migration due to the presence of magmatic fluids is an important element both for the loss of rock strength (destabilizing factor) as well as for the increase of impermeability that promote fluids and water (e.g. meteoric) storage. Once the collapse occurred, the presence of fluids and water, and the clay content as well, can play a fundamental role in the transport and mobility of the avalanche body, enhancing its mobility and transforming the moving mass into a cohesive lahar. Thus, considering all the endogenous and exogenous factors, a volcanic sector collapse and debris avalanche motion depend on an intricate system of factors that promote volcano instability, trigger the collapse, and enhance the avalanche mobility. It is therefore important to understand all the antecedent factors and triggers

Factors Contributing to Volcano Lateral Collapse

of volcano instability since that knowledge can provide critical precursory information that can directly inform hazard warnings. Acknowledgements We acknowledge Marc-André Brideau and Lucia Capra reviewers for their important suggestions and corrections that helped to notably improve the manuscript.

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118 Shea T, van Wyk de Vries B (2010) Collapsing volcanoes: the sleeping giants’ threat. Geol Today 26 (2):72–77 Shea T, van Wyk de Vries B, Pilato M (2008) Emplacement mechanisms of contrasting debris avalanches at Volcán Mombacho (Nicaragua), provided by structural and facies analysis. Bull Volc 70:899-921 Siebert L (1984) Large volcanic debris avalanches: Characteristics of source areas, deposits, and associated eruptions. J Volcan Geoth Res 22:163–197 Siebert L, Roverato M (2020) A historical perspective on lateral collapse and debris avalanches. In: Roverato M, Dufresne A, Procter J (eds). Volcanic dedris avalanches: from collapse to hazard. Advances in volcanology. Springer, Berlin Siebert L, Alvarado GE, Vallance JW, van Wyk de Vries B (2006) Large-volume volcanic edifice failures in Central America and associated hazards. In: Rose WI, Bluth GJS, Carr MJ, Ewert JW, Patino LC, Vallance JW (eds) Volcanic hazards in Central America. Geol Soc Am Spec Pap 412:1–26 Siebert L, Glicken H, Ui T (1987) Volcanic hazards from Bezymianny-and Bandai-type eruptions. Bull Volcan 49(1):435–459 Smith JR, Malahoff A, Shor AN (1999) Submarine geology of the Hilina slump and morpho-structural evolution of Kilauea volcano, Hawaii. J Volcanol Geotherm Res 94(1–4):59–88 Thomas ME, Petford N, Bromhead EN (2004) The effect of internal gas pressurization on volcanic edifice stability: evolution towards a critical state. Terra Nova 16:312–317 Tibaldi A (1995) Morphology of pyroclastic cones and tectonics. J Geophys Res 100(B12):24521–24535 Tibaldi A (1996) Mutual influence of dyking and collapses at Stromboli Volcano, Italy. Geol Soc London Spec Publ 110:55–63 Tibaldi A (2001) Multiple sector collapses at Stromboli volcano, Italy: how they work. Bull Volcan 63(2– 3):112–125 Tibaldi A, Lagmay AMF, Ponomareva VV (2005) Articles effects of basement structural and stratigraphic heritages on volcano behaviour and implications for human activities (the UNESCO/IUGS / IGCP Project 455). Episodes 28(3):158–170 Tinti S, Pagnoni G, Zaniboni F (2006) The landslides and tsunamis of the 30th of December 2002 in Stromboli analysed through numerical simulations. Bull Volcanol 68:462–479 Tommasi P, Baldi P, Chiocci FL, Coltelli M, Marsella M, Pompilio M, Romagnoli C (2005) The landslide sequence induced by the 2002 eruption at Stromboli volcano. In: Landslides: risk analysis and sustainable disaster management. Springer, Berlin, pp 251–258 Tormey D (2010) Managing the effects of accelerated glacial melting on volcanic collapse and debris flows: Planchon-Peteroa volcano. Southern Andes Glob Plan Ch 74(2):82–90

M. Roverato et al. Tuffen H (2010) How will melting of ice affect volcanic hazards in the twenty-first century? Phil Trans R Soc a 368:2535–2558 Ui T (1983) Volcanic dry avalanche deposits—identification and comparison with nonvolcanic debris stream deposits. J Volcan Geoth Res 18(1–4):135–150 Ui T, Yamamoto H, Suzuki-Kamata K (1986) Characterization of Debris Avalanche deposits in Japan. J Volcan Geoth Res 29(1–4):231–243 Urlaub M, Petersen F, Gross F, Bonforte A, Puglisi G, Guglielmino F, Krastel S, Lange D, Kopp H (2018) Gravitational collapse of Mount Etna’s southeastern flank. Sci Adv 4(10):eaat9700. Usami T (1996) Materials for comprehensive list of destructive earthquakes in Japan, 416-1995. Revised and Enlarged Edition. University of Tokyo Press, Tokyo Valderrama P, Roche O, Samaniego P, van Wyk de Vries B, Bernard K, Mariño J, (2016) Dynamic implications of ridges on a debris avalanche deposit at Tutupaca volcano (southern Peru). Bull Volcan 78:14 Villemant B, Komorowski JC, Dessert C, Michel A, Crispi O, Hammouya G, Beauducel F, De Chabalier JB (2014) Evidence for a new shallow magma intrusion at La Soufrière of Guadeloupe (Lesser Antilles). J Vol Geoth Res 285:247–277 Vallance JW, Scott KM (1997) The Osceola mudflow from Mount Rainier: sedimentology and hazard implication of a huge clay-rich debris flow. GSA Bull 109:143–163 Vallance JW, Siebert L, Rose WI, Girón JR, Banks NL (1995) Edifice collapse and related hazards in Guatemala. In: Ida Y, Voight B (eds) Models of magmatic processes and volcanic eruptions. J Volcanol Geotherm Res 66:337–345 Van Bemmelen RW (1949) The geology of Indonesia: The Hague, Government Printing Office, p 732 van Wyk de Vries B (1998) Extension induced by volcanic loading in regional strike-slip zones. Geology 26(11):983–986 van Wyk de Vries B, Davies T (2015) Landslides, debris avalanches, and volcanic gravitational deformation. In: The encyclopedia of volcanoes. Academic Press, pp 665–685 van Wyk de Vries B, Matela RJ (1998) Styles of volcanoinduced deformation: numerical models of substratum flexure, spreading and extrusion. J Vol Geoth Res 81:1–18 van Wyk de Vries B, Merle O (1996) The effect of volcanic constructs on rift fault patterns. Geology 24 (7):643 van Wyk de Vries B, Kerle N, Petley D, (2000) A sectorcollapse forming at Casita volcano. Geology 28:167– 170 van Wyk de Vries B, Self S, Francis PW, Keszthelyi L (2001) A gravitational spreading origin for the Socompa debris avalanche: J Vol Geoth Res 105:225–247

Factors Contributing to Volcano Lateral Collapse Verrucci L, Tommasi P, Boldini D, Graziani A, Rotonda T (2019) Modelling the instability phenomena on the NW flank of Stromboli Volcano (Italy) due to lateral dyke intrusion. J Vol Geoth Res 371:245– 262 Vidal N, Merle O (2000) Reactivation of basement faults beneath Volcanoes: A new model of flank collapse. J Vol Geoth Res 99(1–4):9–26 Villemant B, Komorowski JC, dessert c, Michel A, Crispi O, Hammouya G, Beauducel F, DE Chabalier JB (2014) Evidence for a new shallow magma intrusion at La Soufrière of Guadeloupe (Lesser Antilles) Insights form long-term geochemical monitoring of halogen-rish hydrothermal fluids. J Vol Geoth Res 285:247–277 Voight B (2000) Structural stability of andesite volcanoes and lava domes. Philosophical Transactions of the Royal Society of London. Series A: Math Physic Engin Sci 358(1770):1663–1703 Voight B, Elsworth D (1997) Failure of volcano slopes. Geotechnique 47(1):1–31 Voight B, Glicken H, Janda RJ, Douglass PM (1981) Catastrophic rockslide avalanche of May 18. In The 1980 Eruptions of Mount St. Helens, Washington (1250: 347–377). US Geol Surv Prof Pap Voight B, Janda RJ, Glicken H, Douglass PM (1983) Nature and mechanics of the Mount St Helens rockslide-avalanche of 18 May 1980. Geotechnique 33:243–273 Voight B, Komorowski JC, Norton GE, Belousov AB, Belousova M, Boudon G, Young SR (2002) The 26 December (Boxing Day) 1997 sector collapse and debris avalanche at Soufriere Hills volcano, Montserrat. Geol Soc London Memoirs 21(1):363–407 Voight B, Sousa J (1994) Lessons from Ontake-san: a comparative analysis of debris avalanche dynamics. Eng Geol 38(3–4):261–297 Wang J, Ward SN, Xiao L (2019) Tsunami Squares modeling of landslide generated impulsive waves and its application to the 1792 Unzen-Mayuyama megaslide in Japan. Eng Geol 256:121–137 Watanabe K, Hoshizumi H, Itaya T (1993) K-Ar ages of Unzen Volcano in Kyushu, Japan: with some aspects of geology of Mayu-Yama

119 Watt SFL, Karstens J, Micallef A, Berndt C, Urlaub M, Ray M, Desai A, Sammartini M, Klaucke I, Böttner C, Day S, Downes H, Kühn M, Elger J (2019) From catastrophic collapse to multi-phase deposition: flow transformation, seafloor interaction and triggered eruption following a volcanic-island landslide. Ear Plan Sci Lett 517:135–147 Watt SFL, Karstens J, Berndt C (2020) Volcanic-island lateral collapse and their submarine deposits. In: Roverato M, Dufresne A, Procter J (eds) Volcanic Dedris Avalanches: from collapse to hazard. Advances in volcanology. Springer, Berlin Waythomas CF, Wallace KL (2002) Flank collapse at Mount Wrangell, Alaska, recorded by volcanic massflow deposits in the Copper River lowland. Canad J Earth Sci 39:1257–1279 Waythomas CF, Miller TP, Begét JE (1999) Record of late Holocene debris avalanche and lahar formation at Iliamna Volcano. Alaska. Am Geophys Union Trans 80:F1140 Williams R, Rowley P, Garthwaite MC (2019) Reconstructing the Anak Krakatau flank collapse that caused the December 2018 Indonesian tsunami. Geology 47 (10):973–976 Wooller L, van Wyk de Vries B, Murray JB, Rymer H, Meyer S (2004) Volcano spreading controlled by dipping substrata. Geology 32(7):573-576 Wooller L, van Wyk de Vries B, Cecchi E, Rymer H (2009) Analogue models of the effect of long-term basement fault movement on volcanic edifices. Bull Volc 71(10):1111–31 Zernack AV, Cronin SJ, Bebbington MS, Smith IEM, Price RC, Procter JN (2012) Forecasting catastrophic stratovolcano collapse, a model based on Mt. Taranaki, New Zealand. Geology 40:983–986 Zimbelman DR (1996) Hydrothermal alteration and its influence on volcanic hazards Mount Rainier, Washington, a case history. PhD Thesis, Univ. Colorado, Boulder Zimbelman DR, Watters RJ, Firth IR, Breit GN, CarrascoNuñez G (2004) Stratovolcano stability assessment methods and results from Citlaltepetl, Mexico. Bull Volc 66:66–79

Climatic Influence on Volcanic Landslides Gioachino Roberti, Nicholas J. Roberts, and Catherine Lit

Abstract

Volcanic landslides are controlled by a combination of magmatic, tectonic and surficial processes, the last of which is predominantly influenced by climate. In this chapter, we consider the influences of present and geologically recent climates on the occurrence of volcanic landslides. We begin by summarizing Quaternary climatic variability to illustrate the wide range of conditions and rates of change experienced by modern edifices. A focus on geologically recent volcanoes is prudent because both their morphologies and evidence of the climatic conditions affecting them are typically better preserved; preQuaternary climates similarly affected edifices that are now largely lost from the geomorphic record. We then review the climatic factors that condition and possibly trigger volcanic landslides, the challenges in dating landslides

G. Roberti (&) Minerva Intelligence Inc., 301-850 West Hastings St. Vancouver, Vancouver, BC V6C11, Canada e-mail: [email protected] N. J. Roberts Mineral Resources Tasmania, 30 Gordons Hill Road, Rosny Park, TAS 7018, Australia C. Lit National Institute of Geological Sciences, University of the Philippines Diliman, 1101 Quezon City, Philippines

and climatic changes, and the difficulties in determining triggers of volcanic landslides. Finally, case studies of present and past climate influences on volcanic landslides collected from scientific literature–covering both subaerial and coastal settings–illustrate several key points: edifice collapses were numerous at the end of the last glaciation; current glacial retreat is conditioning volcanic slope failure in some specific settings; shallow landslides in volcanic environments appear to be increasing due to changes in weather extremes; and sea level fluctuation plays a role in volcanic island collapses. As knowledge on climate variability, volcanic, and surficial processes progresses, the understanding of how climate and its changes affect volcanic landslides will further improve.

1

Introduction

Climate influences the stability of volcanic slopes through diverse pathways and over wide ranging timescales. Although climatic drivers also affect non-volcanic slopes, their influences are in many instances amplified by the structural, mineralogical, and geodynamic properties of volcanic landforms. Climate determines temperature as well as type, intensity, duration, and distribution of precipitation, which together affect slope hydrology and hydrogeology. Precipitation and temperature regimes also influence river and

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glacier activity, which in turn control principal aspects of slope stability including incision, stress distribution, and sedimentation. Volcanic slopes are particularly susceptible to these processes, making climatic conditions and their variation a key factor in causing and triggering volcanic landslides. Volcanoes are typically built quickly, by layering of lava flows and pyroclastic deposits. The resulting edifices comprise heterogeneous, sometimes anisotropic material with weak geomechanical properties due to the clastic nature and the rapid emplacement of volcanic rocks and such composition allows some edifices to deform under their own weight. Volcanism commonly increases strain and can produce structural weaknesses through dyke intrusion, volumetric change of magmatic bodies, and contact metamorphism. Additionally, hydrothermal circulation can reduce geologic strength through alteration to or precipitation of weaker minerals. In light of these various volcanic processes, even small climate-driven variation of pore-water pressure changes can alter the edifice stability. Effects of climate on landslide activity in general (Coe 2016; Crozier 2010; Gariano and Guzzetti 2016), and in high mountainous regions in particular (Clague et al. 2012; Huggel 2009; Stoffel and Huggel 2012), have received increasing attention in recent decades, reflecting increasing understanding of climatic systems and their accelerating change. Although various studies consider influences of volcanic eruptions on climate (Gleckler et al. 2006; Miller et al. 2012) and of climate on volcanic eruptions (Jellinek et al. 2004; Praetorius et al. 2016; Rampino et al. 1979; Swindles et al. 2017; Tuffen 2010; Watt et al. 2013), relatively few authors (e.g. Capra 2006; Capra et al. 2013; Deeming et al. 2010; Tormey 2010) focus on climatic influences on volcanic landslide activity. This in part reflects challenges in reliable attribution of volcanic landslides to climatic triggers, which is especially challenging for pre-historic events. Difficulties in demonstrating causal relationships

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stem from (1) dating climate changes, (2) dating landslides, (3) confirming landslide triggers, and (4) proving that the trigger is linked to a climate variation. Here we explore the wide range of climatic influences in conditioning and triggering of volcanic landslides. We first briefly summarize current knowledge of Quaternary climatic fluctuations as well as the resolution of temporal constraints on climate change and landslides. We then consider the effects of climate and its variation on landslides resulting from volcanic slope failures (including edifice and sector collapses), landslides resulting from mobilization of loose material from volcanic slopes (debris flows), and volcanic island landslides. Finally, we present a few modern examples of volcanic landslides affected by climate variation.

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Climatic Variability

Climate characterizes the trend and variability of atmospheric conditions over seasonal and longer timescales. Weather, by contrast, denotes atmospheric conditions at shorter timescales, and is a critical measure of the variability about a longterm trend. As such, the frequency and intensity of weather phenomena are of direct relevance to the characterization of past, present, and future climates, and are critical issues when considering the triggers of volcanic landslides. Climatic conditions, and Earth-surface processes driven by them, differ spatially as well as temporally across the globe. Their type, magnitude, and variability are principally determined by a combination of factors, the most important of which include latitude, elevation, continentality, and ocean–atmosphere circulation. Changes in these conditions have occurred over widely ranging rates, commonly as cyclical alternations, since Earth’s atmosphere first formed. Although we focus on conditions during the Quaternary Period, the process discussed here were undoubtedly also at play in older geologic times.

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The Quaternary Period–since 2.588 Ma (Pillans and Gibbard 2012)—is of particular interest for understanding climatic influence on volcanic slopes for two principle reasons. Firstly, more is known about the Quaternary than earlier periods due to better preservation of its geologic record. Preservation of Pleistocene and Holocene geomorphic records, which is far rarer for older landscapes, further supports Quaternary paleoclimate interpretations. The youth of these landforms, rocks, and sediments generally also supports the application of higher resolution dating methodologies. Secondly, climatic conditions from the Quaternary have modified modern volcanoes during and since their formation, and thus particularly influence contemporary volcanogenic geohazards.

2.1 Quaternary Paleoclimate and Its Drivers As comprehensive review of climate spanning the last *2.6 Ma is beyond the scope of this chapter, we provide only a cursory summary of typical Quaternary conditions and their variability. For a detailed overview of Quaternary paleoclimate and the broad range of techniques utilized in its reconstruction we refer the reader to Walker (2005) and Bradley (2014). Climate variation during the Quaternary can be inferred using a wide range of paleoclimate records and analysis tools. Much of the current knowledge of global patterns comes from deep-sea sediment records of oxygen isotopes, which provide a proxy for storage of water on land, particularly in large ice sheets (e.g. Lisiecki and Raymo 2005; Fig. 1). Chronostratigraphies of these sediment cores are based on a combination of various numerical methods and geomagnetic field reversals. Their isotopic records represent regionally to globally averaged temperatures, and thus do not indicate the full scale of spatial variability at a given time. Ice cores records extending back to the end of the Early Pleistocene (e.g. Jouzel et al. 2007) provide similar isotope-based proxies of paleo-temperature as well as preserved samples of atmosphere composition.

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More localized studies of geologic and geomorphic records provide additional insights of greater spatial and, sometimes, temporal resolution; they indicate details of climatic conditions by way of the consequent biotic (e.g. Yasuhara et al. 2012), chemical (e.g. Singer 1980), and physical (e.g. Macklin et al. 2012) processes affecting Earth’s surface. Correlation of localized records to the global climate record is enabled by a wide range of dating methods (particularly magnetostratigraphy, palynology, tephrochronology, and radiochronologies), although decreasing resolution of many techniques prior the last glacial cycle typically limits confident association with specific positions on the marine isotope curve prior to Marine Isotope Stage (MIS) 2. Utilizing boundary conditions provided by the diverse aforementioned records, numerical climate modelling can be used to examine or infer drivers of climate patterns as well as paleoclimate conditions between locations of detailed records (e.g. Gleckler et al. 2006; Jouzel et al. 2007; Lunt et al. 2008; Pezzi and Cavalcanti 2001). Repeated, long-lived ice ages are a defining feature of the Quaternary Period, during which Earth’s predominant climate pattern has alternated between climates much cooler than at present (glaciations) and those similar to present (interglacials) (Fig. 1). These cycles shifted from *41 to *100-ka periodicity and from symmetric to strongly asymmetric forms near the end of the Early Pleistocene, with longer duration and typically stronger cool peaks since *1 Ma. Regularity of the cycles relates to the interplay between Earth’s orbital parameters (Berger 1992) that determine the amount and distribution of insolation, although the exact pathways of climate forcing remain unclear (Bradley 2014). Consequent Earth-system interactions–namely albedo fluctuations due to snow/ice and dust cover at higher latitudes, variations in atmospheric CO2 concentration, and removal of regolith–modulate the orbital signal and thus influence the variability, duration, and symmetry of a particular glacial cycle (Ellis and Palmer 2016; Willeit et al. 2019). Temperatures also vary over both much longer and much shorter timescales. The cyclic

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alternation between glacial and interglacial periods is superimposed on a longer term cooling trend that initiated with climate deterioration in the latest Pliocene (ca. 3 Ma). The exact cause of this trend are debated, but modelling suggest it relates to decreasing atmospheric CO2 concentration and, to a lesser degree, sea-surface cooling in the equatorial Pacific, uplift of several major mountain systems, and closure of the Panama seaway (Lunt et al. 2008). On much shorter timescales, decadal to internal changes in

ocean (e.g. Lyon and Camargo 2009; Pezzi and Cavalcanti 2001) and atmospheric (e.g. Shaw et al. 2016) circulation influence the distribution of heat and precipitation. Sporadic, geologically instantaneous events perturbed climate on regional or even global scales, including interruption of ocean circulation from rapid draining of inland waterbodies (Teller et al. 2002) and large volcanic eruptions (Gleckler et al. 2006) or a number of closely spaced eruptions (Miller et al. 2012).Variations in Earth’s heat balance

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together with forcing—including from gradual tectonic shaping of marine basins and rapid freshwater debouches—influence ocean–atmosphere circulation and thus the global transfer of heat and moisture.

2.1.1 Glacial Periods The last glaciation approximates conditions over the previous seven glacial cycles, before which cooling and glacier expansion were generally of lesser magnitude and shorter duration (Fig. 1). Gradual, punctuated build-up of extensive glacier ice on the order of 103 ka or more culminated in a maximum just prior to the end of each glaciation. Storage of vast volumes of water in glaciers, combined with thermal contraction of liquid water under lower temperatures, resulted in global sea-level drops of over 100 m (MurrayWallace and Woodroffe 2014), which provides an additional proxy of the global magnitude of cooling. The MIS record suggests that the last glaciation (MIS 2) was globally the most severe (Fig. 1). The largest ice caps generally reached their greatest extent during this time (e.g. Ehlers and Gibbard 2004), although glaciers in some regions were more restricted likely due to moisture limitation or physiographic impediments. Ice caps heavily affected middle-latitude to high-latitude landmasses (>* 45°) and polar waters. High-relief areas were steeply incised, forming high-angle rock slope locally mantled by glacial and proglacial sediments of varying thickness. Adjacent glacier-free areas were cold enough to develop and support permafrost, which in some conditions extended well into subsequent interglacials, including the Holocene. Such permafrost conditions also occurred at lower latitudes where elevation is suitably high. Large proglacial areas were arid due to rain shadows and katabatic winds formed by thick ice sheets. At low latitudes, the main effects would have been large (commonly 100–120 m since ca. 1 Ma) sea level drop and local changes in atmospheric moisture circulation. Marine regression in unglaciated areas caused the emergence of currently partially or even completely submerged areas, including volcanic

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islands. High mountains at low latitude would have also experienced depression of biomes, treeline, and frost line (e.g. Hooghiemstra and Flantua 2019) and, where high enough, local glaciers and even ice caps (e.g. Roberts et al. 2018) repeated over several glacial cycles. Many areas of tropical or sub-temperate rainforest retracted and were in many cases replaced by open savannah (e.g. Webb 1991).

2.1.2 Interglacial Periods Earth’s modern climates—and surface processes —are broadly representative of conditions during past interglacials, including the Holocene, although some temporal variation occurred (Sirocko et al. 2007). Globally, interglacial climates during four of the five last major glacial cycles (Fig. 1) had peaks that were slightly warmer than at present. Earlier interglacials, especially prior to 1 Ma, were milder (Fig. 1). Each interglacial was short lived relative to the glaciations bracketing them. Interglacials since ca. 1 Ma have followed from a rapid warming trend during which ice culminating in the previous glacial maximum rapidly wasted and sea level rapidly rose. Locally such patterns were complicated due to glacioisostatic rebound, hydro-isostatic loading, and independent tectonic adjustments (e.g. Clague and James 2002). Vast quantities of glacially eroded sediment became available for transport due to removal of ice and the absence of vegetation, together with isostatic base-levels lowering (Church and Ryder 1972); in some regions these paraglacial conditions lasted well into post-glacial time, as illustrated by Early Holocene sedimentation patterns. Broad expanses of permafrost adjacent to or replacing previously glaciated terrain waxed and waned with interglacial climatic variations (Reyes et al. 2010; Tesi et al. 2016) Increased chemical weathering in ice-free areas produced weathering profiles of varying depth. At high latitudes these formed slowly and may be immature even after an entire interglacial cycle. At low latitudes, however, chemical alteration of residual soils formed during previous interglacials produced thick, deeply weathered deposits (e. g. Roberts et al. 2017).

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2.2 Near-Future Climate Trends Climatic conditions forecasted for the remainder of the twenty-first century represent intensified version of the interglacial conditions summarized above. However, these conditions are in response to rates of warming unprecedented in the Quaternary. Consequently, ideal analogues for nearfuture conditions do not exist and forecasting changes to atmosphere-ocean circulation as well as related Earth-surface processes includes substantial uncertainty. Climate projections generally predict drier conditions in dry areas, wetter conditions in wet areas, greater precipitation seasonality in most regions, and temperature increase nearly everywhere (IPCC 2014). Principal implications for these climatic changes largely favour failure of volcanic slopes in diverse settings. Such conditions include more severe storms, accelerating global sea level rise, accelerating permafrost degradation, and loss of vegetation cover from increasingly frequent and severe wildfires.

3

Landslide Ages and Uncertainties

Accurate dating of landslides is generally challenging, if even possible. Beyond the uncertainties intrinsic in a given geochronologic technique, difficulties stem from two main sources. Firstly, landslide preservation may be quite low, especially in the case of failures in unconsolidated material or of small magnitude. Secondly, many landslides must be indirectly dated based on their secondary landscape impacts or using bracketing deposits. Consequently, the context of dated materials is of particular importance whether dating landslides in terrestrial (Lang et al. 1999) or marine (Urlaub et al. 2013) settings. In light of the resolutions, error sources, and temporal limits of these various techniques, uncertainties in dating Quaternary landslides are at best sub-annual for the last *102–103 years, and within decades to centuries until 50 ka, and thousands to millions of years for older periods. Many of these applicable

techniques are used for a wide range of Quaternary studies, including dating of the aforementioned paleoclimate records, which thus suffer from many of the same dating uncertainties.

3.1 Dating Terrestrial Volcanic Landslides Diverse geochronological tools are applicable for dating terrestrial landslides, many of which are compressively reviewed by Lang et al. (1999) and more recently by Pánek (2015). The reliability of age constraints improves when multiple complementary dating tools are combined. Such an approach is rare–combining even two independent dating techniques is atypical–but has received more attention in recent decades (Pánek 2015). Although many techniques are applicable to both volcanic and non-volcanic slope failures, the nature of geologic materials common in volcanic terrains may enable additional chronostratigraphic constraints of eruptive units through tephrochronology, fission-track dating, or40K/40Ar and40Ar/39Ar radiometric ages of flows (e.g. Dirksen et al. 2011; Legros et al. 2000). Terrestrial landslides are most commonly dated by indirect dating, by applying radiocarbon techniques (Lang et al. 1999; Pánek 2015), with increasing use of Accelerator Mass Spectrometry (AMS). However, age uncertainties stem from the time lag between landslide deposits and the dated materials (e.g. Friele and Clague 2004), variable timing of plant death (e.g. Evans and Brooks 1991), and variable, often detrital, charcoal sources (e.g. Stoopes and Sheridan 1992). Materials underlying or incorporated into landslide deposits provide only maximum landslide ages, and only until the relatively young (ca. 50 ka) limit of the radiocarbon technique. Dendrochronology provides higher resolution ages and is a popular dating tool (Lang et al. 1999) for more recent landslides. Trees growing on a landslide provide minimum ages due to delayed colonization (e.g. Jakob and Friele 2010) whereas trees buried or entrained by a landslide

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can be dated to provide a maximum age; only trees killed or wounded by the landslide directly date the failure (Salaorni et al. 2017). The issues for radiocarbon and tree-ring dating similarly apply to various geochronologic tools. Furthermore, the common practice of dating sediments deposited in landslide-dammed basins, including radiocarbon, optically stimulated luminescence, magnetostratigraphy, and palynology, yields age estimates that lag an unknown time behind landslide occurrence. Direct dating methods include terrestrial nuclide exposure dating (Cerling and Craig 1994; Dunai 2010; Gosse and Philips 2001) typically of multi-metre-size boulders exposed on the landslide surface or of bare rock at the landslide scar, and 40Ar/39Ar dating of thin frictional melt zones at the base of volcanic landslides (De Blasio and Elverhøi 2008; Lavallée et al. 2012; Legros et al. 2000). However, boulders may have rotated or been exhumed after emplacement, and surfaces may have been exposed prior to the landslide or may have been eroded (Pánek 2015). Such issues lead to ages that either predate or postdate the landslide by an uncertain amount. Frictional heating at the landslide base may have not completely reset the argon clock (cf. Legros et al. 2000). Because of these uncertainties, and due to inherent errors underlying geochronologic approaches, directly dating methods do not necessarily provide more accurate estimates for the timing of past landslides compared to indirect techniques. Additionally, each technique has differing limitations in terms of both resolution and the temporal range over which it is applicable.

3.2 Dating Volcanic Island Landslides Reviews of landslide-dating methods (Lang et al. 1999; Pánek 2015) do not specifically consider failures entering or initiating within the marine domain. This omission partially reflects the fact that largely or entirely marine landslides are much less commonly dated (Urlaub et al. 2013; Whelan and Kelletat 2003), which in turn likely

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reflects difficulties in recognition their submerged morphology and sampling access. For this reason, dating tools for such landslides are considered here in more detail than those of their terrestrial equivalents. Dating efforts usually target large volcanic failures, particularly in the Hawaii, Canary, and Réunion islands (Masson et al. 2002; Whelan and Kelletat 2003) despite the much wider distribution of such landslides (cf. Blahůt et al. 2019), and to a lesser degree especially large non-volcanic failures (e.g. Haflidason et al. 2005). This emphasis on particularly large volcanic failures (e.g. Urlaub et al. 2013) partly reflects the important, but poorly constrained (Masson et al. 2006) secondary hazard posed by waves, as indicated by historic events (Roberts et al. 2014) and modelling of theoretical failures (e.g. Kelfoun et al. 2010; Ward and Day 2001). Due to this secondary effect, the ages of large coastal or submarine failures can also be constrained by dating of associated tsunami deposits (e.g. McMurtry et al. 2004; Paris et al. 2017; Rubin et al. 2000); presence of pumice in the wave deposits provides an addition opportunity for radiometric dating of events as well as an indication that eruptive activity accompanied wave generation and possibly contributed to failure (e.g. Carey et al. 2001; Nishimura 2008; Paris et al. 2017). Volcanic island landslides in marine settings are commonly only loosely dated using a small number of broadly limiting ages. Similar to subaerial settings, 40 K/40Ar (e.g. Labazuy 1996; Masson 1996) and 40Ar/39Ar (e.g. Jicha et al. 2012) radiochronologies of volcanic units below, cross-cut by, or covering landslide scarps. Marine carbonate skeletons (MCS) of molluscs and coral provide several dating options. 14C AMS of MSC of coral macrofossils (e.g. Lipman et al. 1988 using an age from Moore and Fornari 1984) is commonly used (cf. Urlaub et al. 2013). However, the approach is slightly more completed than terrestrial 14C dating as marine reservoir effects must be corrected (Stuiver and Polach 1977).

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Uranium-series disequilibrium dating can date reefs disturbed by landslides and post-failure reef colonization of shallow marine surfaces where coral is present. If applied to molluscs, caution is required because post-mortem geochemical exchange with marine or groundwater sources can introduce significant uncertainty (McLaren and Rowe 1996). Although the potential of both the uranium-series and amino-acid racemisation approaches has been previously recognized in the context of volcanic islands (Whelan and Kelletat 2003), neither is yet typically applied to flank failures. However, oxygen isotope stratigraphy of planktonic foraminifera entrained by landslides can help constrain their ages (e.g. McMurtry et al. 1999) by matching to existing isotope curves. The ages of submarine slope failures can in some cases be better constrained using bracketing deposits since the marine sedimentary record is generally far more continuous than those of most terrestrial settings, enabling the establishment of more consistent chronostratigraphies. The resolution of such control is proportional to the rate of sedimentation, which may be very low in some marine settings. In the absence of knowledge of precise sedimentation rates, the thickness of overlying sediments can help generally estimate the landslides age (e.g. Watts and Masson 1995; Lipman et al. 1988) or at least relative timing of landslides (e.g. Masson et al. 2002). For sediments younger than 50 ka, sedimentation histories are most commonly determined from14C AMS ages on marine shells microfossils (Thomson and Weaver 1994). Chronostratigraphic control is most robust when it includes the sequence both above and below landslides (e.g. Urlaub et al. 2013). If the landslide material is too thick or too stiff to penetrate, however, coring of distal volcaniclastic turbidites may be a viable alternative for obtaining upper and lower limiting ages on failure, provided the turbities can be linked to specific landslide deposits. If cores record suitably long periods they can be dated chronostratigraphically based on paleomagnetism, d18O variation, and excess 230Th profiling (e.g. Hunt et al. 2011; Masson 1996;

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McMurtry et al. 1999), although such approaches are generally best suited to the imprecise dating of pre-Late Pleistocene landslides.

4

Climatic Drivers of Landslides and Field Examples

Slope stability is influenced by numerous exogenous and endogenous factors whose combination over various time scales enables slope failure. Many such factors are affected by climatic conditions and their variation. Which climatic factors affect volcanic slopes and deposits depends on the geographic location of volcanoes (Fig. 2), including whether they are sub-aerial or sub-aqueous, as well as on the volcanoes’ elevation. To conceptualize climatic effects on slopes, we use the ‘factor of safety’ approach (e.g. Crozier 2010) to describe their stability. The factor of safety is a ratio measure of the overall strength of a slope and the net stresses acting on it; when the ratio drops below unity (i.e. net stresses exceed net strength) failure is incipient. Parameters controlling failure can be differentiated as conditioning and triggering factors (Fig. 3). Conditioning factors—including slope geometry, rock-mass fabric, discontinuities, lithologic contrasts, and degree of weathering— determine material conditions such as cohesion, bulk density, failure surface/zone depth and steepness, pore water pressure, and internal friction (Crozier 2010). Triggering factors are events that change these parameters in a short time span, such as intense rainfall, enhanced snowmelt, or ground acceleration, and rapidly push the factor of safety below unity. Climate has long-term and short-term effects on both conditioning factors and triggering factors. In the short term, climatic phenomena are the most common landslide triggers, other than earthquakes (Gariano and Guzzetti 2016). Rainfall and snowmelt are particularly common triggers of landslides in volcanic, as well as nonvolcanic, terrain. Numerous studies demonstrate and explore climatic triggering of landslides and how modern climate change affects it, and the mechanisms generally involve pore water

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Fig. 2 Distribution of volcanos a globally and b by latitude. Data are from the National Centres for Environmental Information’s Volcano Locations Database (https://www.ngdc.noaa.gov/hazard/volcano.shtml) from

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volcanoes are shown by red circles. Plate boundaries: convergent—red; divergent—yellow; and transform— orange

Fig. 3 Long-term and short-term effect of climate change on slope stability. Modified from Huggel et al. (2012)

pressure increases in meta-stable slopes, leading to the failure (e. g. García-Herrera et al. 2005; Christanto et al. 2009; Crozier 2010; Gariano and

Guzzetti 2016a; Segoni et al. 2018). Given it has been accepted that landslide activity is affected by climatic conditions and that its occurrence

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increases during humid periods and deglaciation, landslide clustering has been used as proxy for climate variation (e.g. Borgatti and Soldati 2010; Korup 2012; Trauth et al. 2003; Zerathe et al. 2014). The long-term effects of climate on slope stability are complex. First, climate variation causes fluctuations in mean, maximum, and minimum temperatures as well as changing precipitation types, patterns, magnitudes. These temperature and precipitation changes determine rainfall intensity and frequency of typhoons and storms as well as glacier and river dynamics, slope hydrology, vegetation distribution, and sea level at local to global scales. At the slope scale, variation of rivers discharge influences erosional and sediment transportation potential (Arnell and Gosling 2013), which in turn changes the morphology of slopes and landslide activity. Similarly, glaciers erode, load, and unload slopes changing slope geometries and stress regimes (Grämiger et al. 2017; McColl et al. 2010) and increasing landslide activity following deglaciation (Deline et al. 2015a, b; Evans and Clague 1994; Huggel et al. 2012; Morino et al. 2019). Isostatic depression of the crust by ice sheets and subsequent rebound in response to their melting influences hydrologic base level, leading to increased levels immediately after the post-glacial incision. Post-glacial instability is part of the paraglacial landscape response in previously glaciated and adjacent periglacial environments. For example, glacial sediments in disequilibrium with the fluvial system following glacial retreat increases landslide activity and sediment yield (Mercier et al. 2013) and after that decreases asymptotically with time (Church and Ryder 1972). This phenomenon is pronounced during the immediate post-glacial interval, although it is still occurring in some areas at present as a consequence of modern glacial retreat (Peres and Cancelliere 2018; Wieczorek and Glade 2005). Eustatic sea level fluctuations between glacialinterglacial cycles of up to 130 m (Shackleton 1987) affects stress regimes and porewater pressures of submarine and coastal slopes and determines the location of coastal erosion.

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Additionally, sea level fluctuation can destabilize gas-hydrates that in turn trigger submarine slope failures (Evans et al. 1996). Both the fall and rise of global sea level would affect landslide activity: sea level lowering leads to deep incision of valleys, and steepening of slopes, reducing buttressing; sea level rise, increases pore water pressure in slopes, and increase coast erosion (Day et al. 1999; Keating and McGuire 2004). Several authors (e.g. Goldstrand 1998; Kuijpers et al. 2001; Quidelleur et al. 2008) suggest that global sea level rise and crustal adjustments during and following deglaciation are more efficient in triggering landslides than is climatic deterioration at the onset of a glacial period. This difference is due in part to the comparatively rapid nature of deglaciation at the end-glacial transition. In addition to river and glacier dynamics, temperature and precipitation affect slopes directly through their influences on runoff and soil saturation, wetting and drying cycles (Handwerger et al. 2019), dissolution, weathering (Brooks et al. 2004), permafrost degradation and freeze–thaw cycles (Deline et al. 2014; Draebing et al. 2017, 2016). Colluvium and other sediments on slopes may be in disequilibrium with changing climatic conditions, providing substantial sediment sources for debris flow and landslide activity (Kim et al. 2018; Schmidt and Dehn 2000; Yumul et al. 2011; Christanto et al. 2009; García-Herrera et al. 2005). Finally, climate influences slope stability through the controls that temperature and moisture exert on vegetation distribution. Vegetation can promote slope stability by moderating runoff and anchoring soil masses, thus reducing shallow landslides in the short term (Chirico et al. 2013; Reubens et al. 2007). Conversely, plant roots can deteriorate rock-mass quality, increasing fracture size, facilitating water infiltration, and increasing weathering potential (Drever 1994).

4.1 Volcanic Flank and Edifice Collapse Due to their inherent structure, volcanic edifices may be particularly sensitive to the

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aforementioned climatic conditions. Volcanoes are built quickly in the context of geologic time, rough commonly sporadic emplacement of layered volcanic and volcaniclastic deposits sometime over periods as short as a few days (Pioli et al. 2008). Rapidly emplaced lithologic sequences tend to have heterogeneous and strongly anisotropic geomechanical properties, resulting in stress distributions and alteration potential that favour instability. Moreover, hydrothermal systems commonly alter minerals to types with greatly reduced geomechanical strength (Pola et al. 2012). Many clay minerals formed through hydrothermal alteration and claysize particles are directly deposited by volcanic activity. Clay swelling and shrinking in response to changes in water content further weakens volcanic slopes. Repeated climate-driven wetting and drying drives cyclical expansion and contraction that can over long periods greatly reduce slope strength, both at the surface and at depth. Additionally, tectonism and magmatic intrusion can further weaken volcanic edifices, which also deform under their own weight (Cecchi et al. 2004). When volcanoes fail, a large part of the volcanic edifice may be involved in the landslide, generating some of the largest subaerial mass movement on earth (van Wyk de Vries and Delcamp 2014). During such events, the presence of fines enhances the mobility of volcanic landslides, which have longer runout compared to non-volcanic landslides of similar volume (van Wyk de Vries and Delcamp 2014). In this context, even small climate-driven changes on volcanic slopes can have devastating effects. Considering paleo-climate change effect on volcanic edifice stability, Capra (2006) noticed an absence of volcanic collapse in the geologic record during glaciations in both the Northern and Southern hemispheres, in stark contrast to common volcanic collapses during subsequent rapid deglaciation in the Late Pleistocene and Early Holocene. Examples of post-glacial collapses include Antuco volcano (Chile), Volcán de Colima and Jocotitlán (Mexico), Taacapa (Chile) and Egmont (New Zealand). Capra (2006) acknowledges the difficulty in determining whether the triggers for most such events

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were climatic or seismic. Capra et al. (2013) focuses specifically on Late Pleistocene Mexican volcano collapses. They demonstrated the importance of moist climates and consequent elevated pore water pressures, both as a trigger of those collapses and as a contributor to their transformation into long-runout, volcanic debris flows. Moisture content records from lake sediments and speleothems provide a proxy for winter and summer precipitation to support the climatic trigger hypothesis. Wet climates also influenced the collapse of Etna volcano that formed the Bove Valley. Deeming et al. (2010) observed a correlation between the moist period at 7500 BP in Sicily and the subsequent collapse. The authors argued in favour of a possible magmatic component in the collapse: dyke emplacement in a water saturated flank caused pore water pressure to increase, triggering the failure. While this trigger mechanism is not definitively determined, the climatic component in determining the flank water saturation is. Tormey (2010) discusses the Planchon-Peteroa volcanic collapse at the Late Pleistocene/Early Holocene deglaciation in Chile and explores the effect of present-day glacial retreat from ice-capped volcanoes. The author proposed a framework for glacio-volcanic hazard and risk assessment, as the human population is now more at risk compared to the Late Pleistocene age. In Canada, high landslide activity following Pleistocene deglaciation is documented both for Garibaldi and Mount Meager volcanoes. The western flank of Garibaldi volcano collapsed at the end of the Pleistocene glaciation, generating the large Cheekye fan (Mathews 1952). Friele et al. (1999) and Friele and Clague (2009) studied the sedimentary structure of the fan, concluding that about 80% of the sediments of the fan were deposited within 4000 years of Pleistocene deglaciation. In contrast, Mount Meager volcano, has continuously been the site of very large landslides (106–108 m3) and significant debris flow activity through the Holocene (Friele and Clague 2009; Friele et al. 2008) until present day, thus also providing a great example of present-day glacial retreat related landslides.

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Mount Meager has been argued to be the most unstable mountain in Canada (Friele et al. 2008) and has been the site of at least three recent (1975, 1998, and 2010) large volcanic landslides triggered by ice and snow melting during intense summer heat waves (Roberti et al. 2018). Access to the area is now restricted following climatic thresholds which consider temperature and precipitation (Friele 2012). Many other large slopes, ripe for catastrophic failure, have been identified at Mount Meager volcano (Roberti 2018). Figure 4 shows the West flank of Plinth Peak, in the Mount Meager volcanic complex. At Mount Meager the complex interaction between volcanic, gravitational and glacial processes is evident: a large volcanic flank formed by weakened volcanic materials is actively deforming under its own weight; at the same time a glacier is retreating from the toe, and hot volcanic gases are piercing through the ice, forming ice caves (Roberti 2018). Similarly, Iliamna Volcano in Alaska is a glaciated volcano that generates rock and ice avalanches with volumes of 106–107 m3 every few years (Huggel et al. 2007). At Iliamna, the avalanches are mainly related to a glacier collapsing from a steep slope, involving some hydrothermally altered rocks. The thermal perturbation of the ice by the geothermal heat flux is believed to be the main cause of these avalanches, but observations at Iliamna lead the authors to discuss the role of temperature changes in cold environments as a destabilizing factor. The transition from cold based to warm

based ice reduces the friction between the bed and the ice, eventually leading to glacier collapse. This observation needs further investigation but opens discussion on the effects of temperature changes in cold environments. In tropical settings, a present-day climatetriggered volcanic landslide example is provided by the 1998 Casita volcano event (Kerle and van Wyk de Vries 2001; Scott et al. 2005; van Wyk de Vries et al. 2000). Intense rainfall brought by hurricane Mitch in October 1998 caused the collapse of about 200,000 m3 of the hydrothermally altered and sagging volcanic edifice. Cumulative three-day rainfall just prior to collapse was 1217 mm, nearly four times the monthly average (328 mm) (Kerle and van Wyk de Vries 2001). The collapse transformed into a long-runout volcanic debris flow that buried two towns, killing *2500 people.

Fig. 4 Panoramic photograph of the large moving slope of Plinth Peak, in the Mount Meager volcanic complex, Canada. Upper and lower slope deformation rates, ice

limits, and fumarole are shown; note the knob of yellowaltered rocks at the centre of photograph. Helicopter and people provide scale

4.2 Shallow Volcanic Landslides and Debris Flows Volcanic edifices are also prone to frequent smaller scale failures including rockfalls, shallow landslides, and debris flows. Small scale failure can directly transition into debris flows, or can accumulate debris readily mobilized by rainfall (e.g. Paguican et al. 2009) or rapid snowmelt (e.g. Decaulne et al. 2005), and generate debris flow, especially because volcanic rock materials are commonly already very weakened by

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hydrothermal processes, many small, lowmobility rock failures (e.g. falls or slumps) that would otherwise remain intact instead disintegrate to produce more mobile failures that in the presence of water will also be more likely to produce debris flows. Mountain streams can generate debris flows when water starts to mobilize sediments. Depending on the rock types underlying the stream watershed, basins can be classified as transport-limited or climate-limited (Bovis and Jakob 1999). Transport-limited drainages are controlled by the availability of sediments in the channel and have strong underlying rocks: after an event, it takes some time to erode the strong rocks and to accumulate debris in the channel again. On the other hand, climate-limited drainages have an unlimited supply of sediments; these occur in highly erodible rocks, where after reaching some precipitation threshold, a debris flow can always be generated from direct erosion of sediments and weak rocks in the watershed. The high erosion potential of freshly emplaced pyroclastic deposits or hydrothermally altered rocks make volcanoes especially prone to climate-limited debris flow, both in temperate and tropical volcanoes (Bovis and Jakob 1999; Rodolfo and Arguden 1991). For example, the 1984 Mayon eruption was followed by many years of intense debris flow activity, with most events occurring in the years immediately following the eruption (Rodolfo and Arguden 1991). The 1991 Pinatubo eruption was also followed by intense debris flow activity (Pierson et al. 1996; Umbal and Rodolfo 1996). In the months following the eruption, between July and November 1991, about 185  106 m3 of material was transported on the southwest sector of Pinatubo, affecting an area of 46 km2 and remobilizing about 14% of the erupted material (Umbal and Rodolfo 1996). Also recently exposed, poorly vegetated surfaces in glaciated settings provide abundant, readily mobilized sediments. In Canada, high sediment yield is documented for Garibaldi, Cayley, and Mount Meager volcanoes (Bovis and Jakob 1999). Especially at Mount Meager, Friele and Clague (2009) calculated denudation rates of

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3000 m3 km−2 year−1 which is 50-fold the average rate of the non-volcanic mountains in British Columbia. The volcano is also affected by the Little Ice Age retreat (Holm et al. 2004; Roberti 2018) that presently conditions landslide and debris flow activity (Fig. 5). Increased frequency and magnitude of storms bring stronger winds and more rainfall, increasing the likelihood of volcanic debris flows from tropical to temperate volcanoes. In the Philippines, for example, tropical typhoons have increased in frequency and intensity since 1971 (Cinco et al. 2016) and have triggered significant recent events. In November 2006, Typhoon Reming exceeded any present storm on record and brought rainfall with an intensity of 47.5 mm/h for one and a half days on the southwestern flank of Mayon Volcano, triggering debris flows that caused 1266 fatalities (Paguican et al. 2009). On glaciated volcanoes, change in frequency, magnitude, and type (snow vs rain) of storms adds to the sediments recently exposed by glacial retreat. Example of this phenomenon is provided by volcanoes of the Cascade Range (Burns et al. 2015), when in 2006 Madden Julian Oscillation brought 50 mm of precipitation over four days, triggering large debris flows. This system, colloquially called the ‘Pineapple Express’, came earlier that year and all the precipitation occurred as rain, greatly increasing runoff.

4.3 Volcanic Islands Landslides In addition to the aforementioned climate factors affecting subaerial edifices, volcanic islands are affected by marine processes including water temperature fluctuation, gas hydrate destabilization, sea level variation. Extensive debris avalanche deposits on the seafloor adjacent to most coastal and volcanic islands serve as evidence for their past large-scale failure (Blahut et al. 2019; Holcomb and Searle 1991). These deposits have been studied by multiple authors at Hawaii (e.g. Garcia 1996; Moore and Clague 1992; Morgan et al. 2003), Piton de la Fournaise (e.g. Labazuy 1996; Lénat et al. 1989; Oehler et al. 2008),

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Fig. 5 Google Earth view of retreating glaciers at Mount Meager volcano. Glacier limits in 1981 and 2019 glacier outlines (glacier is partly debris covered), as well as some

scarps are indicated. Note abundant sediments in the main valley and the gullies on the valley sides. The 1981 glacier position is from Roberti (2018)

Martinique (e.g. Brunet et al. 2015; Friant et al. 2015), Stromboli (Kokelaar and Romagnoli 1995; Romagnoli et al. 2009a, b), Augustine Island (Begét and Kienle 1992), the Canary Islands (Hunt et al. 2018, 2013; Krastel et al. 2001; Masson 1996; Watts and Masson 1995), and others. Comprehensive volcanic island edifice stability is presented by Keating and McGuire (2000) and in Watt et al. (2020)—this volume. The climatic component of volcanic island landslides has also been reviewed by Keating and McGuire (2004). Generally, those authors conclude that landslide activity occurs at volcanic islands during post-glacial periods. At this time, the sea level rise is accompanied by wet and warm climate on land. The wet climate can saturate the volcanic edifice, and the slopes weakened by erosion during previous low stands are now more susceptible to failure both by surficial processes and by earthquake or dyke intrusion (Day et al. 1999; Keating and McGuire 2004). For example, Ablay and Hurlimann (2000) discussed landslides at Tenerife as potentially related to sea level lowstand at 180 ka (MIS 6) and 560 ka (MIS 14). In contrast, Boulesteix et al. (2013) link the collapses of La Orotava at Tenerife and La Cumbre Nueva at La Palma to the rise of sea level and/or to humid conditions at

the onset of the interglacial stage (MIS 13) at about 550 ka (Fig. 6). At Gran Canaria, Lomoschitz et al. (2002) discuss increased landslide activity during wet and humid interglacial climates from the Miocene to the Pleistocene. These precise correlations stem from good constraints on the ages of the collapses, as these are followed by datable eruptions. The same correlation is not possible at all volcanic islands; for example, for Hawaiian volcanoes, ages are not well constrained and the relation between climatic variation and landslides is not definitive (Keating and McGuire 2004).

5

Conclusions

Climate exerts important influences on volcanic landslides. Improved records of Quaternary climate change provide key insights on the atmospheric phenomena and consequent Earth-surface process responsible for the shaping, modifying, and weakening of existing volcanoes and their deposits. These processes can condition and ultimately trigger volcanic mass movements at a range of scales and through multiple mechanisms as demonstrated by the literature review of climatic drivers of volcanic landslides. The role

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Fig. 6 The large Orotava collapse at Tenerife, Canary Islands, which could be related to MIS 13 sea level rise and wet conditions (Boulesteix et al. 2013)

climate plays depends on prevailing long-term conditions and short-term extremes at the volcano site, type and composition, and mechanisms of landsliding. Examination of three contrasting volcanic landslide types–edifice collapses, shallow landslides and island collapses–highlights many of these controlling factors. Due to inherent uncertainties in dating both landslides and climatic fluctuations, demonstrating climate as the landslide failure trigger is more difficult than demonstrating its role in determining the longterm controlling factors. Additionally, characterizing controls and potential triggers of past failures becomes more challenging with increasing age since paleoenvironmental reconstruction, including both the nature and timing of past environments, also becomes more difficult. Despite difficulties in assessing the precise roles of climatic conditions and the precise timing of failures, volcanic landslides are undoubtedly affected by the same processes that affect non-volcanic landslides. However, the response of volcanic slopes is enhanced by the dynamic interaction between surficial, magmatic, and tectonic processes. Climate has been shown to be a significant conditioning factor of volcanic landslides but has only been demonstrated to be their trigger in relatively few cases. Confirming whether a volcanic landslide was climatically induced necessitates parsing out all the possible

magmatic, tectonic and surficial components, which is challenging if even possible. However, there is general agreement that clustered volcanic landslide activity is typically related to climatic change. Current substantial gaps in the understanding of past climatic variability and its influences on volcanic slope stability will continue to narrow as our understanding of both climatic and volcanic systems improves and as dating techniques are expanded and refined. Acknowledgements We would like to thank Lisa Borgatti and Costanza Morino for a critical review of this paper. Eventual errors or omissions are purely the responsibility of the authors.

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Climatic Influence on Volcanic Landslides van Wyk de Vries B, Kerle N, Petley D (2000) Sector collapse forming at Casita volcano, Nicaragua. Geology 28(2):167–170 Walker M (2005) Quaternary dating methods. Wiley, Chichester, 304 p Ward SN, Day S (2001) Cumbre Vieja Volcano-Potential collapse and tsunami at La Palma, Canary Islands. Geophys Res Lett 28:3397–3400 Watt SFL, Pyle DM, Mather TA (2013) The volcanic response to deglaciation: evidence from glaciated arcs and a reassessment of global eruption records. EarthSci Rev 122:77–102 Watts AB, Masson DG (1995) A giant landslide on the north flank of Tenerife, Canary Islands. J Geophys Res 100(B12):24487–24498. doi: 10.1029/95JB02630 Webb SD (1991) Ecogeography and the great American interchange. Paleobiology 17:266–280 Whelan F, Kelletat D (2003) Submarine slides on volcanic islands—a source for mega-tsunamis in the quaternary. Prog Phys Geogr 27:198–216 Wieczorek GF, Glade T (2005) Climatic factors influencing occurrence of debris flows. In: Hungr O, Jackob M

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Volcanic Debris Avalanche Transport and Emplacement Mechanisms Engielle M. R. Paguican, Matteo Roverato, and Hidetsugu Yoshida

Abstract

or within the avalanche body. Multiple shear zones include progressive fragmentation within the avalanching mass, resulting in pockets of shear and slip. We present case studies for each model and hypotheses for the formation of flowbands on the deposit surface. Processes involved during emplacement include disintegration, dynamic fragmentation, and matrix injection. Near the base, bulldozing and incorporation of substrata change the composition and behaviour of the VDA. In extreme cases, VDAs transform into lahars if sufficient water is available for entrainment. Post-emplacement, lahars can also happen, e.g., through debris dewatering, loading of saturated substrata or in the case of landslide dam failure. VDA also create secondary slides when deflected by topographic barriers or when the margins are oversteepened.

Field observations of volcanic debris avalanche (VDA) morphology, sedimentology, and structural features have inspired several hypotheses on their dynamic behaviour. These include plug flow, translational slide, and sliding along multiple shear zones, none of which involve large-scale turbulence during transport. The plug flow model shows normal gradation in the plug, and reverse grading in the laminar boundary layers. During translational sliding, spreading of the mass is accommodated by listric normal faults that flatten into a main sliding plane at the base of

E. M. R. Paguican (&) School of Geography, Environment and Earth Science, University of the South Pacific, Laucala Campus, Suva, Fiji e-mail: [email protected] E. M. R. Paguican College of Engineering and Geosciences, Caraga State University, Ampayon, Butuan City, Philippines M. Roverato Department of Earth Sciences, University of Geneva, Geneva, Switzerland M. Roverato School of Earth Science, Energy and Environment, Yachay Tech University, Urcuqui, Ecuador H. Yoshida School of Arts and Letters, Meiji University, Tokyo, Japan

1

Introduction

Volcano flank destabilisation is a long-term process, induced by continual forces such as tectonic activity (Lagmay et al. 2000; Vidal and Merle 2000), gravitational spreading (Fig. 1a, b; Borgia et al. 1992; van Wyk de Vries and Francis 1997); internal growth by magmatic intrusion (Fig. 1c; Donnadieu and Merle 1998; Tibaldi 2001); and weakening by hydrothermal alteration (Fig. 1d; van Wyk de Vries and Francis 1997;

© Springer Nature Switzerland AG 2021 M. Roverato et al. (eds.), Volcanic Debris Avalanches, Advances in Volcanology, https://doi.org/10.1007/978-3-030-57411-6_7

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Fig. 1 Pre-avalanche gravitational deformation conditions by a spreading (Borgia and van Wyk de Vries 2003) and b sagging, and a weak core due to c magma intrusion or d hydrothermal alteration (Cecchi et al. 2005)

Reid et al. 2001). Catastrophic failure of a destabilised volcano flank, however, is a geologically instantaneous event that can generate large volcanic debris avalanches (VDAs), triggered by either earthquakes (Montalto et al. 1996), magmatic intrusion (Voight et al. 1983; Elsworth and Voight 1996), meteoric events (van Wyk de Vries et al. 2000), a combination of these events, or even without obvious trigger (cf. Roverato et al. 2020, this volume). During growth, a stratovolcano can have multiple lateral collapses (Belousov et al. 1999; Vidal and Merle 2000; Tibaldi 2001, 2005; Roverato et al. 2011; Paguican et al. 2012; Siebert and Roverato 2020, this volume; Dufresne et al. 2020, this volume).

The causes and triggering mechanisms of VDAs are relatively well constrained (Siebert et al. 1987; McGuire 1996; Voight and Elsworth 1997; van Wyk de Vries et al. 2001; Roverato et al. 2020, this volume), whereas their emplacement mechanisms are still poorly understood (Francis and Wells 1988; Siebe et al. 1992; Cleary and Campbell 1993; Legros et al. 2002). Given that VDAs are a major volcanic hazard and can be highly destructive (Leyrit 2000), we need to better understand their transport and emplacement mechanisms to be able to predict runout and assess potential hazards in volcanic areas. We define transport mechanisms as the flow regimes during the spreading of the

Volcanic Debris Avalanche Transport and Emplacement Mechanisms

avalanche body and emplacement mechanisms as those that happen as the avalanche wanes and towards the final arrest of the sliding mass. Direct observations of the transport and emplacement mechanisms of VDAs are extremely rare but forensic evidence preserved in the sediments and in the morphological features provides clues. This chapter first presents the morphological features of volcanic debris avalanche deposits (VDADs) that are used for interpreting the transport and emplacement mechanisms of VDAs. Then, we present the important processes that occur during transport, the flow regimes and emplacement models that could operate in combination, we also discuss debris avalanche related granular flows, and flow transitions during transportation (syntransportation) and after emplacement (syndeposition) of VDADs. This chapter uses the terms syn-transportation and syn-depositional lahars and landsliding to differentiate these events when associated with a debris avalanche from those that are not. For a detailed discussion of the causes and triggers of volcano lateral collapse and a comprehensive description of VDAD facies and general terminology, please refer to Roverato et al. (2020, this volume), Bernard et al. (2020, this volume), and Dufresne et al. (2020, this volume).

2

Morphological Features

VDADs are easily recognised by their irregular, hummocky topography, characterized by small depressions and mounds, and longitudinal and transverse ridges that stand out of the surrounding topography (Siebert 1984; Francis and Wells 1988; Glicken 1996; Belousov et al. 1999; Shea and van Wyk de Vries 2008; Dufresne and Davies 2009; Andrade and van wyk de Vries 2010). Studying the surface features and internal architecture of these deposits, we can find evidence of the transport and emplacement mechanisms of these avalanches as well as the processes that modify the avalanche during and after their final arrest.

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2.1 Hummocks, Ridges, and Flowbands Hummocks are circular or elongated surface features (Fig. 2a; Siebert 1984; Francis and Wells 1988; Glicken 1996; Shea and van Wyk de Vries 2008; Dufresne and Davies 2009). Larger hummocks called torevas are found in the proximal area of some VDADs, within or at the foot of the failure scar (Reiche 1937; Voight et al. 1981, 1983; Wadge et al. 1995; Belousov et al. 1999; van Wyk de Vries et al. 2001; Kelfoun and Druitt 2005; Shea and van Wyk de Vries 2008; Thompson 2009; Andrade and van Wyk de Vries 2010; Paguican 2012). Torevas exhibit backward rotation toward the source area. Hummocks and torevas are the morphological expressions of brittle deformation, and are related to the strain and deformation regimes within the VDAD. They are the remains of initial failure blocks that slide, tilt, and rotate down the slope (e.g., Paguican et al. 2014; Dufresne and Geertsema 2020). Hummocks start with fracturing and normal faulting of edifice blocks, then extension of the spreading debris forces the upper brittle layer to split and drop adjacent blocks, forming horst and graben structures. Ridges are more elongated, well-defined, topographic highs, and can be parallel or perpendicular to spreading directions (Valderrama et al. 2018). Ridges that extend the entire flow length separated by parallel bands of narrow trenches have been called flowbands at Tutupaca, Peru (Valderrama et al. 2018) and elsewhere, striations at Shiveluch, Russia (Belousov et al. 1999), and herringbone at Lastarria, Chile and Argentina (Fig. 2b, c; Naranjo and Francis 1987). At Lastarria, V-shaped ridges point uphill. In distally raised VDADs, compressional structures formed during avalanche deceleration such as those seen in Chaos Jumbles, Carlson, Martinez Mountain, Blackhawk, Lastarria, Aucanquilcha, Ollagüe, Tetivicha, Llullaillaco, Parinacota, Socompa and Olympus Mons are referred as transverse ridges (Belousov et al., 1999; Shea and van Wyk de Vries 2008; Dufresne 2009; Dufresne and Davies 2009).

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Fig. 2 Morphological features in volcanic debris avalanche deposits: a Hummocks at Iriga, Philippines (Paguican et al. 2014); b Elongated ridges and flowbands described as striations at Shiveluch, Russia (Belousov et al. 1999); c herringbone flowband structures at

Lastarria, Chile (Naranjo and Francis 1987); d fingers and e lobes at Tutupaca, Peru and in analogue experiments (Valderrama et al. 2016, 2018); f Aerial photograph and g sketch of faults at Llullaillaco in Chile and Argentina (Shea and van Wyk de Vries 2008)

Ridges form in analogue experiments with a certain percentage of fines that is vibrated, expelling coarse particles to the top and causing particle segregation (Cassie et al. 1988;

Pouliquen et al. 1997; Pouliquen and Vallance 1999; Malloggi et al. 2008; Gray and Kokelaar 2010; Johnson et al. 2012). This particle segregation resulted in three distinct morphologies:

Volcanic Debris Avalanche Transport and Emplacement Mechanisms

joint fingers, separated fingers, and poorly developed fingers or lobes (Fig. 2d–e; Valderrama et al. 2018). Dufresne and Davies (2009), however, suggests that formation of surface mounds appears to be an intrinsic process of free-surface granular flows and their appearance as ridges, flowbands, or aligned hummocks depends on the frictional behaviour of the material, the emplacement velocity and direction, the emplacement geometry, and the influence of the substrate on the flow dynamics. Analogue experiments by Shea and van Wyk de Vries (2008) suggest that ridged and faulted VDAD result from homogeneously sized or homogeneously competent rock fractions such as in Socompa VDAD where 90% of its total volume is made up of fine substratum, whereas differences in the cohesion, competence, size, and lithology of the source materials such as the alternating sand and plaster in analogue volcanoes or alternating different lithologies in volcanic environments like in Mombacho, Parinacota, Ollagüe and Tetivicha, produce hummocky VDAD topography (van Wyk de Vries et al. 2001). Additional discussion on hummock and ridges formation is in Sects. 4.2 and 5. Dufresne and Davies (2009) presented a continuum of hummock morphologies based on the length to height (L/H) ratio: hummocks have L/H ratios less than or equal to 10, ridges have ratios greater than 10, and flow bands have ratios in the hundreds.

2.2 Faults and Folds VDADs often exhibit well-developed faults (Fig. 2a, f–g; Shea and van Wyk de Vries 2008). Differential movements within the VDA generate strike–slip faults. Spreading (i.e., lateral and longitudinal extension) is accommodated by normal and strike–slip faults, whereas upon deceleration, thrust faulting, folding, and thickening of distal debris may dominate (Shea and van Wyk de Vries 2008; Dufresne 2009; Paguican et al. 2014). The faults and topographic profile of a VDAD can be used to investigate whether and where within the mass the avalanche

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had a dominantly compressional or extensional regime. Predominantly compressional avalanches typically have shorter runouts and cover smaller areas than extensional-dominated avalanches of similar volumes (Shea and van Wyk de Vries 2008; Paguican et al. 2012, 2014).

3

Processes Acting During VDA Emplacement

VDAs are granular flows wherein the collapsing material disaggregates through processes like particle breakage, dynamic fragmentation, shear concentration, or abrasion. During transport, processes act within the body, or at the surface or near the base of VDAD regardless of the dominant emplacement model described in the following sections. They are important factors in sustaining the dispersion of the main body of the avalanche and for explaining the high mobility of volcanic avalanches.

3.1 Disintegration, Dynamic Fragmentation, and Mechanical Fluidization Within the inner part of a granular mass column, inter-particle collision leads to progressive grain size reduction during the entire transport phase of the avalanche. This process results in an increasing proportion of matrix relative to blocks with distance travelled (Perinotto et al. 2015). Perinotto et al. (2015) state that dynamic disintegration occurs continuously for larger particles throughout the transport, as long as topographic confinement remains sufficient to maintain an effective frictional or shear stress, intense syntransport fragmentation, and collisional fracturing and crushing prevail. After fracturing, inflation affects the jigsaw clasts, and progressive dispersion of sub-particles within the matrix results in some degree of dispersion of the granular flow. Fragmented particles that escape further crushing are smoothed and rounded by frictional abrasion, also reducing their size. The

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initial kinetic energy of the avalanche during the initial acceleration of the avalanche body and lithostatic pressure variations caused by the loading and unloading of grains above the fragmenting particles, impacts the shearing and fragmentation within the granular mass (Tavares and King 1998). More recently, Zhang and McSaveney (2017) presented that in rock avalanches, normal stress and shear strain are the most significant factors controlling fragmentation. Further fracturing and dilatancy of the granular mobile mass comes from the elastic energy released during dynamic fragmentation (Davies et al. 2019a, b) and vibrations induced by their movement (Melosh 1979, 1987; Collins and Melosh 2003). Dynamic (basal) fragmentation is a dispersive force that can keep granular material in a dilated state (Davies et al. 2010). It is the response of intact brittle rock to rapid strain under a confining stress. During this process, rocks deform elastically, storing elastic strain energy from the general shearing motion around it until its local strength is exceeded. It then breaks and fragments along fracture surfaces, and a substantial proportion of its stored elastic strain energy is released as the kinetic energy of the fragments moving away from the original centre of mass. This generates a basal shear stress that increases the motion of the avalanche, and that is a function of the flow thickness and the intact rock's strength. This was the basis for a constant basal retarding stress in numerical runout model that simulated the low basal shear resistance, spreading behaviour, and runout distance of the Socompa VDAD and the deflection by topographic barriers that caused a secondary slide (Davies et al. 2010). Melosh (2015) suggests the fluctuating pressures of strong vibrations created acoustic waves within the mass that fluidise homogeneous avalanching debris. Individual clasts push tightly against each other because of the overburden pressure, and then suddenly become free to slip due to pore pressure fluctuations within the normally clast-supported matrix. This causes laminar flow within the VDAD that does not mix distinct rock units, despite very large strains

E. M. R. Paguican et al.

incurred during transport. Mechanical fluidisation explains the high rates of shearing at the base of debris avalanches. In this process, high energy input from, for example, a severe earthquake that lasts several minutes or sufficient seismic and acoustic excitation through dynamic fragmentation causes high impulsive contact pressures between individual grains, so that they become statistically separated and the mass dilates (e.g., Davies 1982). This reduces internal resistance to shear stress resulting in a locally high dilation and reduction of internal friction. Lubrication and fluidisation are important mechanisms at VDA transport and emplacement (Hughes 1970; Howard 1973; Scheidegger 1973; Hsü 1975; McSaveney 1978; Melosh 1979; Campbell et al. 1995; Davies 1982; Anders et al., 2000; Collins and Melosh 2003; Pudasaini and Miller 2013).

3.2 Substrate Entrainment and Deformation At or near the base, shearing is accommodated by structures such as normal and reverse faults with associated fractures affecting the substratum, sheared blocks and matrix that have penetrated the substratum, block stretching along the contact, laterally continuous undulations of the contact, lamination in the substratum, and erosional stripping of mechanically weak and deformable VDAD layers (Fig. 3a–d; Friele and Clague 2004; Bernard et al. 2008; Caballero and Capra, 2011; Paguican et al. 2012; Tost et al. 2014; Roverato et al. 2018). In Chimborazo (Table 1), substratum-derived content was estimated to be 50–70 vol.% of the mixed facies, giving a volume of 10–14 vol.% of the entire 1.2–1.6 km3 volume of the VDAD (Bernard et al. 2008). This component reached up to 35 vol.% in the distal area of the Pungarehu VDAD at Taranaki (Palmer and Neall 1991). This proves substratum incorporated during transport. As the proportion of finer material increases with distance, rheology can change during travel. When an avalanche incorporates a significant volume of the substrata

Volcanic Debris Avalanche Transport and Emplacement Mechanisms

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Fig. 3 a Debris avalanche transport and emplacement mechanism model, involving the incorporation of the substratum by injection and bulk fluidisation [modified after Bernard et al. (2008)]. b Deformed lacustrine megablock at the base of El Zaguán debris avalanche deposit in Nevado de Toluca, Mexico [see also Caballero

and Capra (2011)]. c Lava domains imbedded in the yellowish interclast matrix and sharp basal contact with the older succession; and d two different ripped-up clasts eroded from the substrate during transport in Cubilche debris avalanche deposit in northern Andes, Ecuador [see also Roverato et al. (2018)]

(*14 vol.%), it can develop a fine lubricating basal layer depending on the type of substrata incorporated, i.e., its material properties and composition, which can result in enhanced mobility. These basal structures are easier to observe towards distal areas because of the

thinner deposits and erosion have acted on these parts more efficiently. Erosion and entrainment through rip-up clasts of soil and tephra have been observed in the older Iriga, Ruapehu, and Pylon Peak VDADs (Friele and Clague 2004; Paguican et al. 2012; Tost et al. 2014), to a limited extent

Translational Slide

Plug flow

Model

Tongueshaped

Chimborazo

Lobate, tongue-shaped

Lobate, distally raised

Lobate, symmetrical, fan-shaped

Irregular, proximally raised

Fan-shaped, distally raised

Fan-shaped, asymmetric, proximallyraised

Socompa

Mombacho “Las Isletas”

Mombacho “El Crater”

Parinacota

Older Iriga

Younger Iriga

Kaida

Iwasegawa

Shape

VDAD

Hummocks; ridges

Faults, hummocks, torevas

Faults, hummocks, torevas, longitudinal and transverse ridges

hummocks

Faults, hummocks

Faults, longitudinal and transverse ridges

hummocks

Hummocks and ridges

Surface features

Partly confined

Unconfined

Confined

Partly confined at north

Unconfined at south

Initially unconfined, later confined

Confined

Confined

Confined

Confinement

Mainly flat but channelled by surrounding mountains

Along river valleys

Along river valleys

On a basin

Runout path topography

1300 (precollapse)

6350 asl

1345 asl (current)

3000 (summit and basin)

3063 (current)

1178 (current)

3300

Summit height (m)

2

>6

1.8

1.2

36

>0.3

*0.1

>11

Volume (km3)

70

118

140

49.5

57

500

280

12

16

22

12.4

12

40

46

12.4

35

Length (km)

(continued)

Area (km2)

Table 1 General characteristics and geometric data of VDAD case studies used as basis for the transport and emplacement mechanisms of VDAs (Naranjo and Francis 1987; Belousov et al. 1999; Takarada et al. 1999; Clavero et al. 2002; Bernard et al. 2008; Kelfoun et al. 2008; Shea et al. 2008; Shea and van Wyk de Vries 2008; Andrade and van Wyk de Vries 2010; Davies et al. 2010; Paguican et al. 2012; Roverato et al. 2015; Valderama et al. 2016, 2018)

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Granular flow

Some hummocks Hummocks; ridges

Tutupaca

Hummocks; ridges

Faults, longitudinal and transverse ridges

Shiveluch

Lastarria

Almost fan shaped

Fan slightly outward

Pungarehu

Surface features

Multiple shear zones

Shape

VDAD

Model

Table 1 (continued)

Confined

Unconfined

On a plain and confined by old glacial valleys for most of its path

Unconfined

Confinement

Into a gentle, locally dissected, south-sloping plain

Runout path topography

1000

>4000 (precollapse)

Summit height (m)

0.091

40

3% (Scott and Janda 1987). In contrast, time-delayed reworking processes most commonly result in the generation of sanddominated debris flows and hyperconcentrated flows (Major et al. 2005; Procter et al. 2009; Zernack et al. 2009; Zernack 2020, this volume)

197

or transitions between them (cf. Pierson and Scott 1985). Scott et al. (2001) summarized the textural changes associated with the progressive transformation from landslide to cohesive debris flow (cf. their Fig. 15). The transformation is accompanied by changes in megaclast/clast-to-matrix ratios (decreasing), grain size distribution (developing from coarse-skewed to bimodal), and the fracturing and disaggregation of megaclasts (from relatively coherent clasts to domains of clasts and matrix). The four stages of transformation defined by Scott et al. (2001) are best reflected in the texture of the deposit: from the megaclast-dominated stage-1 with locally closed work to stage-2, which is “trimodal with megaclasts (openwork dispersed in bimodal debris flow matrix)” to (stage-3) “scattered megaclasts dispersed in debris flow matrix”, and finally to the “bimodal, with modes separated by granule (2–4 mm) fraction” of stage-4. Sedimentological and textural studies have helped distinguish reworked avalanche deposits (Mt. Ruapehu, Keigler et al. 2011; Nevado de Colima, Capra et al. 2002) from lahar deposits (Guisaugon, Catane et al. 2008; Citlaltepetl, Carrasco-Nuñez et al.1993; Mt. Rainier, Scott et al. 1995), hyperconcentrated flow deposits (Mt. St-Helens, Mt. Rainier, and Mt. Ruapehu), and streamflow deposits in the distal zone (Mt. Rainier, Vallance and Scott 1997; Mt. Ruapehu, Keigler et al. 2011; Tost et al. 2014). It is crucial to recognize the potential of debris avalanches to directly transform into cohesive debris flows as this has implications for hazard assessment of distal areas. The change in flow behaviour due to high initial contents of watersaturated or clay-rich material and/or subsequent incorporation significantly increases the total travel distance, with reports of highly mobile debris flows exceeding debris avalanche runout by almost 100 km (Table 4). Distinguishing these dynamic processes from post-emplacement reworking is an important task in the assessment

198

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Table 4 Examples of debris avalanches that transformed into debris flow during transport/emplacement and the increase in total travel distance Volcano

V (Mm3)

Ltotal (km)

Length percent DF

References

Mt. Rainier, USA

2000– 2500

120

98

Scott et al. (1995), Vallance and Scott (1997)

Nevado de Toluca, Mexico

2800

75

27

Capra and Macías (2000)

Casita, Nicaragua

1.6–2

30

92

Kerle and van Wyk de Vries (2001), Kerle (2002)

Pico de Orizaba, Mexico

200,000

95

21

Capra et al. (2002), Carrasco-Nuñez et al. (1993)

Cerro Rabicano, Chile

15

57

70

Hauser (2002)

Mt. Cayley (1984), Canada

0.9

6.1

43

Evans et al. (2007)

of debris flow dynamics, runout predictions, and hazard assessment.

5.1 Direct Transformation of VDAs into Cohesive Debris Flows Three lithofacies that define distinct mapping units are recognized within the relatively unconfined VDADs from Mt. Taranaki (Neall 1979; Palmer et al. 1991; Alloway et al. 2005; Zernack et al. 2009, 2011). Axial-A and axial-B lithofacies (Table 1) correspond to the debris avalanche phase, whereas the marginal lithofacies is attributed to the debris flow phase. These lithofacies developed in response to changes in the nature of the flow as the VDA travelled away from source (Fig. 9a), and they differ considerably from valley-confined units described elsewhere (e.g. Siebe et al. 1992; Richards and Villeneuve 2001; Clavero et al. 2002; Tost et al. 2014). For example, hummocks are less prominent in confined VDADs, megaclasts are less common, the matrix-to-clast-ratio may be comparatively low, and shearing plays a larger role than in unconfined VDADs (cf. Tost et al. 2014, their Table 1). Axial-A facies forms near-source lobes that exhibit a hummocky surface with closely-spaced, large mounds up to 50 m high and with basal diameters up to 500 m. The internal fabric consists of brecciated, self-supporting megaclasts

and less (120 km (Stoopes and Sheridan 1992). However, it was later suggested that the deposit represents several events related to volcano collapse and subsequent processes (Capra and Macías 2002; Capra et al. 2002). The initial large, water-saturated VDAD with an estimated volume of 7 km3 had a runout distance of 45 km, where it came to a hold at a physiographic barrier, temporarily damming the Naranjo River (Capra and Macías 2002). Rapid

A. Dufresne et al.

reworking of the VDAD surface produced a secondary lahar that inundated the southern plain and entered the Salado River drainage. Subsequent successive failure of the dam remobilized *1.5 km3 of VDAD material in a series of debris flows. These might have merged to form an exceptionally large, catastrophic flow that reached the Pacific coast more than 90 km from the volcano (Capra and Macías 2002). The estimated deposit volume of 10 km3 indicates that the flow bulked up to six times its initial size due to entrainment of underlying sediments and bedrock along the flow path.

6

Conclusions

In the past four decades since the eruption and lateral collapse of Mount. St. Helens volcano, much progress has been made towards fully understanding generation, transport and emplacement mechanisms as well as hazards from volcanic debris avalanches. Advances in analytical methods, and the rigorous interpretation of the vast sedimentological evidence assist in: • Identifying and reconstructing past events— closing gaps in frequency analyses and completing regional assessment of mass wasting processes. • Focusing research questions, which should guide sampling techniques (e.g. questions regarding fragmentation processes in different domains should be based on facies understanding to guide grain size analyses). • Understanding the processes involved in matrix generation and dynamic deformation of the moving avalanche mass by compiling solid field evidence. This evidence collection will help validate or refute existing emplacement hypotheses and guide future (numerical and theoretical) investigations. • Pinpointing the location and identifying the processes of transformation into highly mobile debris flows—an important aspect in hazard assessment since transformation during VDA transport and emplacement will increase the total area affected during the event,

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whereas post-emplacement debris flow initiation might pose a ‘hidden’ hazard that requires risk mitigation to consider longer timescales.

volcano, Kamchatka, Russia. Bull Volcanol 61:324– 342 Bernard B (2008) Etude des dépôts d’avalanches de débris volcaniques: analyse sédimentologique d'exemples naturels et identification des mécanismes de mise en place. Ph.D. dissertation. Université Blaise Pascal, Clermont-Ferrand, France (in French) Bernard B, van Wyk de Vries B, Leyrit H (2009) Distinguishing volcanic debris avalanche deposits from their reworked products: the Perrier sequence (French Massif Central). Bull Volc 71(9):1041 Bernard B, Takarada S, Andrade SD, Dufresne A (2020) Terminology and strategy to describe volcanic landslides and debris avalanches. In: Roverato M, Dufresne A, Procter J (eds) Volcanic debris avalanches: from collapse to hazard. Springer book series advances in volcanology Bernard K (2015) Quelques aspects sédimentaires des avalanches de débris volcaniques. Ph.D. Thesis, Univ. Clermont-Auvergne, France (unpub., in French). Available at :. Bernard K, van Wyk de Vries B (2017) Volcanic avalanche fault zone with pseudotachylite and gouge in French Massif Central. J Volcanol Geotherm Res 347:112–135 Bernard K, Thouret J-C, van Wyk de Vries B (2017) Emplacement and transformations of volcanic debris avalanches—a case study at El Misti volcano, Peru. J Volcanol Geotherm Res 340:68–91 Bernard K, van Wyk de Vries B, Thouret J.-C. (2019) Fault textures in volcanic debris-avalanche deposits and transformations into lahars: the Pichu Pichu thrust lobes in south Peru compared to worldwide avalanche deposits. J Volcanol Geotherm Res 371:116–136 Beuselinck L, Govers G, Poesen J, Degraer G, Froyen L (1998) Grain-size distribution by laser diffractometry: comparison with the sieve-pipette method. CATENA 32:193–208 Billi A (2005) Grain size distribution and thickness of breccia and gouge zones from thin (< 1 m) strike-slip fault cores in limestone. J Struct Geol 27(10):1823– 1837 Blott SJ, Pye K (2001) GRADISTAT: a grain size distribution and statistics package for the analysis of unconsolidated sediments. Earth Surf Proc Landforms 26(11):1237–1248 Borgia A, van Wyk de Vries, B (2003) The volcanotectonic evolution of Concepción, Nicaragua. Bull Volc 65:248 Boule M (1896) Le Cantal miocène. Bull Serv Carte Géol Fr 8(54):213–248 Brideau M-A, Procter JN (2005) Discontinuity orientation in jigsaw clasts from volcanic debris avalanche deposits and implications for emplacement mechanism. GEOQuébec, 8p Brideau MA, Procter JN (2015) Discontinuity orientation in jigsaw clasts from volcanic debris avalanche deposits and implications for emplacement mechanism. GeoQuébec 2015:20–23

This review also highlights the large number of interpretations that are ambiguous or heavily debated. Part of the problem is terminology— differences in nomenclature might point to factual differences in VDA transport and emplacement conditions. In other cases, inconsistent terminology simply causes confusion. This underscores the need to carefully select and justify the terms used in future publications. Acknowledgements We highly appreciate the constructive feedback of Lizeth Caballero and Sergio Salinas who kindly reviewed this chapter and helped improve the clarity and structure with their thoughtful comments.

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Volcanic Debris-Avalanche Deposits in the Context of Volcaniclastic Ring Plain Successions—A Case Study from Mt. Taranaki Anke V. Zernack

Abstract

Volcaniclastic successions represent valuable archives that hold a detailed record of volcanic and other landscape-shaping events and often provide the only way to reconstruct the long-term volcanic history of a region. The ability to accurately interpret the origin, transport and emplacement processes of such deposits is thus crucial to better understand the nature, magnitude and frequency of future volcanic and secondary hazards, including the potential for catastrophic edifice failure. Recognising the sedimentary characteristics of different lithofacies associations in modern ring plain successions also helps evaluate the processes that shaped ancient successions, which often lack a wider depositional context. Continuous coastal erosion of the tectonically uplifted Taranaki Peninsula, New Zealand, has exposed an almost complete stratigraphic record of medial-distal ring plain successions. These unique coastal cross-sections make Mt. Taranaki an ideal case study to assess typical lithofacies associations and sedimentary processes occurring in an unconfined ring plain depositional system around a long-lived

A. V. Zernack (&) School of Agriculture and Environment, Massey University, Private Bag 11 222, Palmerston North, New Zealand e-mail: [email protected]

andesite stratovolcano. This chapter presents the sedimentary characteristics of the wide spectrum of volcaniclastic and reworked sedimentary facies observed at Mt. Taranaki and the frequency of volcanic mass flows. Factors influencing ring plain accumulation and landscape evolution are then put into global context and discussed in the light of modern and ancient volcaniclastic successions elsewhere.

1

Introduction

Volcaniclastic sediments have been described in detail at volcanoes worldwide (e.g. Cas and Wright 1987; Manville et al. 2009a). Previous work often concentrated on the sedimentary characteristics of the exposed deposits in order to infer their origin, possible trigger mechanisms, mode of transport and depositional conditions. This is aimed towards developing a better general understanding of the flow dynamics and behaviour of volcanic mass flows. Reconstruction of short periods of the volcanic history, typically within the younger, more easily accessible record, and the characterisation of individual events are typically used as the basis for future hazard assessments (e.g. Siebert et al. 1995; Scott et al. 1997; Belousov et al. 1999; Waythomas 1999; Waythomas and Miller 1999; Reid et al. 2001). Only a few studies have described in a holistic sense the setting, detailed

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accumulation history and overall evolution of volcaniclastic successions in the surroundings of long-lived stratovolcanoes that form volcaniclastic ring plains or aprons (Palmer and Neall 1991; Smith 1991a; Cronin et al. 1996; Davidson and De Silva 2000; Donoghue and Neall 2001; Borgia and van Wyk de Vries 2003; Alloway et al. 2005; Zernack et al. 2011). The term volcaniclastic “ring plain” was introduced from studies in New Zealand for “flat or nearly flat land in a nearly circular or annular area surrounding Mount Egmont” (now commonly referred to as Mt. Taranaki; Morgan and Gibson 1927). Subsequently the term was applied more generally to the low-relief alluvial plains and fans surrounding the volcanoes of the central North Island, i.e. Mt. Ruapehu and Mt. Tongariro (Palmer et al. 1993; Cronin et al. 1996; Cronin and Neall 1997; Lecointre et al. 1998; Donoghue and Neall 2001). Smith (1987a, b) used the term volcaniclastic apron instead of ring plain to describe “relatively thin (200 ka volcanic history of Mt. Taranaki. Initially, four debris-avalanche deposits were mapped in west and south Taranaki (Neall 1979), while another three units were later identified to the north-east and south-east of the volcano (Alloway et al. 2005). Subsequent detailed mapping of the ring plain succession exposed in south-western coastal cliff sections revealed numerous previously unknown debris-avalanche deposits that are buried under younger sediments and thus do not exhibit a mappable surface expression (Zernack et al. 2009, 2011). At least 14 widespread debris-avalanche deposits are now recognised within the ring plain record (Table 1), as well as several cohesive (clay-rich) debris-flow deposits that likely represent the distal runout of smaller debris avalanches (Zernack et al. 2011). The minimum run-out distance for debris avalanches is marked by the present-day coastline of the Taranaki Peninsula and generally exceeds 25 km to the west, 45 km to the south and 39 km to the north. The northern deposits can be traced for at least another 6 km offshore (Alloway et al. 2005). Volumes range from small deposits (7.5 km3)

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Table 1 Composite stratigraphic overview of volcanic debris avalanche (VDA) and cohesive debris-flow (DF) deposits and other key marker beds of the Mt. Taranaki ring plain succession as discussed in the text

a

Neall (1979), bZernack et al. (2011), cVandergoes et al. (2013), dAlloway et al. (1995), eDanišik et al. (2012), fTinkler (2013), gMcGlone et al. (1984), hPillans (1983), iAlloway et al. (2005)

14

C yrs BP refers to conventional radiocarbon years before AD1950 (of radiocarbon dated contexts) whereas age (ka) means 103 years before present and includes both, calibrated radiocarbon dates BP and other geohistorical dates derived from the rock record (e.g. chronostratigraphic estimates)

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deposited by debris avalanches, with deposit thickness of 2.5 to >16 m in medial coastal cross-sections (Neall 1979; Neall et al. 1986; Ui et al. 1986; Palmer and Neall 1991; Palmer et al. 1991; Alloway et al. 2005; Procter et al. 2009; Zernack et al. 2009, 2011). Debris avalanches were sourced from all sectors of the volcano and inundated all parts of its ring plain, excluding the north-northwest where the flows were deflected around the eroded remnant of the neighbouring Pouakai edifice and the higher surfaces of the Pouakai ring plain (Fig. 3). The resulting voluminous deposits are the main landscape-forming element of the Taranaki ring plain and form broad fans of hummocky terrain (cf. Glicken 1991, 1996)

around the volcano, variously covered by tephra, loess, and lahar deposits. Classic sedimentary features include jigsaw-cracked and shattered clasts, megaclasts, and large rip-up clasts (Figs. 4 and 5; Neall 1979; Palmer and Neall 1991; Palmer et al. 1991; Alloway et al. 2005; Procter et al. 2009; Zernack et al. 2009, 2011; Roverato et al. 2015). Due to their unconfined nature and emplacement on low-relief terrain, Taranaki debrisavalanche deposits developed distinct surface geometries and lithofacies distributions that differ from valley-confined examples (Palmer et al. 1991; Dufresne et al. 2020—this volume; Procter et al. 2020—this volume). Proximal deposits exhibit a brecciated, clast-supported fabric with

Fig. 3 Map showing the dispersal, age and estimated volume of Mt. Taranaki VDA deposits and the Maitahi VDAD from Pouakai Volcano. The youngest (Opua, Motumate, Pungarehu and Ngaere FM) and two older units (Motunui and Okawa Fm) were mapped based on their surface expression (Neall 1979; Alloway et al. 2005), whereas the other VDADs were identified and

mapped in coastal cross-sections (Zernack et al. 2009, 2011). Their extrapolated lateral extent is shown as transparent markers in order of emplacement with solid fills representing preserved exposure in cliff sections. Offshore extent of DAD along the northern coast from Alloway et al. (2005)

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Fig. 4 Characteristic sedimentary features of Mt. Taranaki debris-avalanche deposits. a Close to source, Mt. Taranaki DADs are characterised by different domains with clast-supported fabric representing still intact stratigraphy of the source deposits (1 m-long tape measure for scale). b These shatter to various degrees during debris-avalanche initiation, transport and emplacement producing characteristic jigsaw cracks. c Continuous fragmentation during transport results in an increase in

matrix and a decrease in clast abundance and size (arrows point to jigsaw-fractured clasts) as well as d decreasing size and abundance of megaclasts, typically consisting of brecciated and/or shattered portions of source material (tape measure is 1 m long). e Towards the margins, the deposits are matrix-dominated containing smaller clasts and rip-up clasts (arrows) and f occasional small brecciated clasts. Hammer for scale, handle c. 30 cm long

little sandy matrix (Fig. 4a, b) and a higher density of large mounds along the main dispersal axes that reduce in spatial density and size laterally and with increasing distance from source (Procter et al. 2020—this volume). Medial

deposits display an increasing matrixcomponent, abundant megaclasts, jigsawcracked and shattered clasts as well as rip-up clasts (Figs. 4c, d and 5; Neall 1979; Palmer and Neall 1991; Palmer et al. 1991; Alloway et al.

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Fig. 5 Medial to distal Mt. Taranaki DADs in coastal cliffs contain various types of secondary components that were picked up along the flow path during transport. a The most common types consist of soil fragments or peat with/without intercalated tephra beds (pencil for scale is 15 cm long); b pieces of wood with large tree logs being rare (hammer for scale); c fragments of underlying volcaniclastics, such as debris-flow and hyperconcentrated deposits d that can be of considerable size. e Clasts

made of underlying sandstone and Tertiary mudstone (hammer for scale) are limited to the older debrisavalanche deposits as these were emplaced when the basement was still exposed in some areas, which were subsequently buried under the developing volcaniclastic ring plain. f This type of rip-up clast is very common in the Maitahi FM from Pouakai Volcano and typically strongly deformed

2005; Procter et al. 2009; Zernack et al. 2009). Towards the margins, the deposits grade into thinner, matrix-supported cohesive debris-flow deposits that are characterised by scattered low mounds, a progressive decrease in primary clasts

and an increase in secondary components (Figs. 4e, f and 5). These gradual, consistent facies transitions are best displayed by the Pungarehu Formation, the most voluminous known collapse at Mt. Taranaki, and the youngest

Volcanic Debris-Avalanche Deposits in the Context …

example, the Opua Formation (Neall 1979; Ui et al. 1986; Roverato et al. 2015; Procter et al. 2009, 2020—this volume). The observed longitudinal and lateral transformation of volcanic debris avalanches into long-runout cohesive debris flows has been mainly attributed to the incorporation of watersaturated sediments and soils (e.g. Palmer et al. 1991; Vallance and Scott 1997; Scott et al. 2001). At Mt. Taranaki, fine-grained andic coverbeds and soils/paleosols on the edifice and ring plain contain high proportions of ferrihydrite and allophane, a short-range-order clay mineral that forms by rapid weathering of andesitic ash under the prevailing humid-temperate climate conditions of the region (Neall 1976; Alloway et al. 2005; Procter et al. 2009). During edifice failure, this fine-grained allophane-rich material is incorporated in the generated debris avalanche, both as a primary component derived from the collapsed edifice strata and from eroded ring plain sediments during transport. The known thixotropic properties of allophane and its high capacity for water storage most likely promoted the transformation and resulting extremely high mobility of Mt. Taranaki debris avalanches (Neall 1976; Palmer et al. 1991; Alloway et al. 2005; Zernack et al. 2009). Debris-avalanche deposits are the most useful and reliable stratigraphic marker units for correlation of complex volcaniclastic sequences in Taranaki and elsewhere, due to their distinct sedimentary characteristics, considerable thickness and wide, continuous lateral distribution. More detail on the spatial distribution of all marker beds of the southwestern ring plain, including a cross-section of the coastal cliffs, can be found in Zernack et al. (2011) while the distribution of northern and eastern debrisavalanche deposits has been described by Alloway et al. (2005).

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3.2 Lahar (Debris-Flow and HyperconcentratedFlow) and Fluvial Deposits While individual debris-avalanche units are the most prominent, voluminous and widespread stratigraphic markers in the Taranaki ring plain succession, other types of volcanic mass flow deposits are much more abundant (Neall 1979; Neall et al. 1986; Palmer and Neall 1991). The wide variety of lahar initiation mechanisms results in a broad range of deposits with diverse textures and fabrics (Fig. 6). Their composition, grain size and thickness depend on the type, size and origin of the flow as well as on the depositional environment (Fisher and Schmincke 1984; Scott 1988a, b; Smith and Lowe 1991; Scott et al. 1995b; Cronin and Neall 1997; Vallance 2000; Major et al. 2005). The deposits from debris flows are recognised by their massive, poorly sorted texture, coarse grain-size, matrixsupport of clasts and often inverse grading, which reflect the high sediment concentration, particle interaction, high yield strength, buoyancy and laminar flow of the transport medium (Fig. 7). Hyperconcentrated-flow deposits are better sorted and finer grained, show faint stratification, clast support, normal, inverse or no grading indicating less yield strength and buoyancy and evidence of more turbulent flow behaviour (Fig. 6a–f). Normal streamflow deposits are produced by fully turbulent flow resulting in better sorting and distinct horizontal bedding to cross-stratification (Fig. 6g, h). Lahar deposits can show several facies that reflect the lateral and longitudinal changes in flow dynamics. The progressive downstream transformation from debris flow to hyperconcentrated flow produces deposits that show a gradational vertical transition from a basal hyperconcentrated-flow deposit upward to a

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debris-flow unit (Pierson and Scott 1985; Cronin et al. 1999). With increasing dilution, the basal horizontally bedded, poorly sorted, clastsupported hyperconcentrated-flow layer thickens, while the overlying massive, unbedded, coarse, very poorly sorted, matrix-supported debris-flow portion thins (Scott 1988a; Scott et al. 1995b; Cronin et al. 1999; Zernack et al. 2017). With increasing distance from source, the horizontal stratification becomes more distinct and is eventually replaced by wavy crossstratification as the flow transforms to normal streamflow. Some lahars are vertically stratified into a coarse, sediment-rich channel flow and an overlying dilute, finer grained surface layer, which results in accumulation of near-channel, wedge-shaped debris-flow deposits that are laterally equivalent to overbank hyperconcentratedflow deposits (Fig. 7; Cronin et al. 2000). Clasts within lahar deposits can be primary, mostly comprising angular to subangular volcanic rocks from the source region, or secondary angular to rounded clasts that were picked up along the lahar path (Major and Scott 1988; Scott 1988a, b; Pierson et al. 1990; Major et al. 2005). The clast assemblage can be monolithologic but is more commonly polylithologic, with bimodal grain-size distributions (Vallance 2000). Vesicles commonly found in the matrix result from entrapment of air bubbles. Other components include wood fragments, casts of tree fragments and charcoal. Based on their sedimentary characteristics and lateral facies changes, non-cohesive volcanic mass-flow deposits in Taranaki were classified as channelised debris-flow and related overbank deposits, sheet-like (floodplain) hyperconcentrated-flow deposits, transitional hyperconcentrated flow/normal streamflow deposits and fluvial deposits (Zernack et al. 2009). Channelised, non-cohesive debris-flow deposits are very coarse, poorly sorted, clastsupported with little sandy matrix and typically unstratified (Fig. 7a, d). They grade laterally into thinner, better sorted, faintly bedded overbank deposits that consist of fine pebbly sands (Fig. 7b, c). Distinct from the lahar overbank deposits are tabular hyperconcentrated-flow

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deposits that can be traced up to 2.5 km in lateral exposure (c.f. Figs. 6a and 7b). They typically contain pebbly clasts in a sandy matrix and show a variety of different characteristics ranging from poor to moderate sorting, massive to bedded, and graded to non-graded (Fig. 6). Marginal features that indicate transition to normal stream flow deposits include strongly developed horizontal to wavy bedding or low-angle crossbedding, abundant lenses of cross-bedded fine sands and rounded pumice lapilli and a lenticular geometry with very erosive basal contacts, often forming steep, overlapping channels (Fig. 6e). Differences in sedimentary characteristics are the result of different flow source regions, flow dynamics and emplacement conditions as well as diversity in paleo-depositional environments. Vertical facies transitions of individual units reflect variations in flow regime and sediment load with time, while lateral facies variations are attributed to distance from source, dispersal axes or channel geometry and capacity (Procter et al. 2009; Zernack et al. 2009). Fluvially reworked volcaniclastic deposits are highly localised and consist of alternating layers made of moderately to poorly sorted, sand- to boulder-sized, sub-rounded to rounded volcanic clasts and laminated, well-sorted, fine to coarse sands (Fig. 6g, h). The deposits show low angle cross-stratification, prominent scour-fill crossbedding and erosive contacts to the underlying sediments.

3.3 Primary Volcanic Deposits (Lavas, Tephras and Pyroclastic Flow Deposits) Primary volcanic deposits are mostly restricted to the modern edifice and proximal parts of the younger ring plain due to their often-limited dispersal/run-out as well as subsequent rapid weathering, reworking or burial. At Mt. Taranaki, frequent effusive activity and domebuilding episodes generated lava flows and dome-collapse block-and-ash-flows (BAFs), which built up the present-day edifice (Fig. 8; Neall 1979; Neall et al. 1986; Platz et al. 2007).

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224 b Fig. 6 Hyperconcentrated-flow deposits dominate the

Mt. Taranaki volcaniclastic succession, which is characterised by deposits produced by a wide spectrum of sediment-water flows. a Sediment-rich hyperconcentrated flows emplaced coarse and reverse to normally graded or massive and ungraded units, the latter showing transitions to debris-flow deposits. b More dilute flows produced bedded, fine-grained hyperconcentrated-flow deposits, some exhibiting pumice “trains” and discontinuous beds of aligned coarser particles. Some HFDs are almost monolithologic and can be c pumice and scoria-rich or d dominated by dense andesite clasts with occasional breadcrust bombs, indicating syneruptive origin or generation shortly after eruptive activity. e Hyperconcentrated

While the oldest recognised lava flows are 130 to 28 ka) are preserved in peat and carbon-rich sediments along the northern coast (Alloway et al. 2005).

3.4 Aeolian Sands, Marine Terraces, Peats and Paleosols At least two spatially separate sets of crossbedded aeolian sands occur in the Taranaki ring plain succession. They are characterised by alternating beds of dark grey, heavy mineral-rich sands and coarser yellow to brownish, pumicerich sands (Fig. 9a). The so-called Punehu Sands, consisting of up to 12 m-thick wellsorted, cross-stratified sands, are exposed over a >15 km-wide stretch of the south-western coastline (Fig. 9b; Zernack et al. 2009, 2011). In some areas, single or multiple hyperconcentrated sheet-flow deposits and localised fluvial sediments are intercalated, while primary tephra layers are absent, suggesting rapid saltation and redeposition of tephras as lenses of coarser rounded pumice lapilli. The presence of numerous thin peat beds, iron-stained weathering horizons and weakly developed tephric soils in the tops of individual dune sets indicate accumulation over a long period of time during cool as well as mild climates. Their stratigraphic position and similarity to present-day near-shore dune fields suggest that they formed in a relatively undisturbed near-coastal environment following MISS 5a, one of the higher sea level stands of the last Interglacial, giving an

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Fig. 7 Geometries of channelised lahar deposits in the Mt. Taranaki ring plain succession. a A typical wide channel sequence consists of a central channel area filled by coarse debris-flow units characterised by a clastsupported fabric and large boulders that grade into finergrained overbank deposits near the channel margin, overlain by finer-grained sheet-like hyperconcentratedflow deposits. b Abrupt facies changes from coarse debris-flow deposits filling a steep-sided channel to finergrained overbank facies. c These facies changes are more

gradational in wide, gently sloping channels often found within the larger river systems. d Major channel in the large river systems (RS) can extend laterally for up to 60 m, here the Lizzie Bell RS, where the channel fill exhibits wavy contacts to the partially eroded underlying Otakeho DAD. e Some parts of the large river systems are characterised by smaller channels cut by hyperconcentrated flows and smaller debris flows (arrows), here the Opunake RS overlain by the Pungarehu and Opua DADs

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Fig. 8 a The modern-day edifice of Mt. Taranaki seen from the West with satellite cone Fanthams Peak (FP), The Beehive domes and steep lava bluffs that mark the amphitheatre created by the 7.5 ka Opua flank failure; marginal Opua VDA mounds in the foreground. b Arrows point to the older lava ridge that marks the extent of the Opua amphitheatre, which has been almost completely infilled with lavas, pyroclastic material and the since developed Fanthams Peak. c Over the past 1,000 years, BAFs and lahars have mostly been directed to the NW due to the breached summit crater configuration, adding to the development of a large volcaniclastic apron that has been building up since 12 ka. d View towards the

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breached crater area, which is surrounded by steep lava flow cliffs and the remnants of the summit dome. While lava flows make up the largest proportion of the modern edifice, the flanks also comprise significant volumes of lahar and BAF deposits with intercalated pyroclastic fall and flow beds. e The upper Pyramid Stream is dominated by lahar deposits comprising reworked BAF material (cliffs are c. 30 m high). f To the East, fallout sequences are more prominent such as at Curtis Ridge where BAF deposits are overlain by a 28 ka, pre-Ngaere in the eastern sector. A minimum of 14 widespread debris avalanche and a number of cohesive debris flow deposits, interpreted to represent the runout of smaller debris avalanches, are recognised in the >200 ka volcanic record of Mt. Taranaki (Zernack et al. 2011). This suggests that a major edifice failure occurred on average every 14,000 years with an apparent increase in frequency since 40 ka to one event every

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6700 years, although this might simply indicate more complete preservation and exposure of the younger volcanic record. This semiregular recurrence pattern (one failure every 2.5– 25 kyrs) is the result of a long-term steady magma supply rate and thus growth process punctuated by random triggering of edifice failures. Collapse frequency and size appear to be controlled by the nature of the triggering event and the precondition of the edifice at the time (Zernack et al. 2012a; Zernack and Procter 2020 —this volume). While the current annual collapse probability is *0.00018, with the most likely collapse being a small one (1 Ma) record offshore Montserrat, Lesser Antilles. Geochem Geophys Geosyst 17:2591–2611 Crandell DR (1989) Gigantic debris avalanche of Pleistocene age from ancestral Mount Shasta volcano, California, and debris-avalanche hazard zonation. US Geol Surv Bull 1861:1–32 Crutchley G, Karstens J, Berndt C, Talling PJ, Watt SFL, Vardy ME, Huhnerbach V, Urlaub M, Sarkar S, Klaeschen D et al (2013) Insights into the emplacement dynamics of volcanic landslides from highresolution 3D seismic data acquired offshore Montserrat, Lesser Antilles. Mar Geol 335:1–15 Day SJ (1996) Hydrothermal pore fluid pressure and the stability of porous, permeable volcanoes. Geol Soc Spec Pub 110:77–93 Day SJ, Heleno da Silva SIN, Fonseca JFBD (1999) A past giant lateral collapse and present-day flank instability of Fogo, Cape Verde Islands. J Volcanol Geotherm Res 94:191–218 Day S, Llanes P, Silver E, Hoffmann G, Ward S, Driscoll N (2015) Submarine landslide deposits of the historical lateral collapse of Ritter Island, Papua New Guinea. Mar Pet Geol 67:419–438 Deplus C, Le Friant A, Boudon G, Komorowski JC, Villemant B, Harford C, Ségoufin J, Cheminée JL (2001) Submarine evidence for large-scale debris avalanches in the Lesser Antilles arc. Earth Planet Sci Lett 192:145–157 Dufresne A, Davies T (2009) Longitudinal ridges in mass movement deposits. Geomorphology 105:171–181 Elsworth D, Day SJ (1999) Flank collapse triggered by intrusion: the Canarian and Cape Verde Archipelagoes. J Volcanol Geotherm Res 94:323–340 Frey-Martínez J, Cartwright J, James D (2006) Frontally confined versus frontally emergent submarine landslides: a 3D seismic characterisation. Mar Pet Geol 23:585–604 Garcia MO (1996) Turbidites from slope failure on Hawaiian volcanoes. Geol Soc Lond Spec Publ 110:281–294 Garcia MO, Hull DM (1994) Turbidites from giant Hawaiian landslides: results from ocean drilling program site 842. Geology 22:159–162 Gee MJR, Watts AB, Masson DG, Mitchell NC (2001) Landslides and the evolution of El Hierro in the Canary Islands. Mar Geol 177:271–293 Germa A, Quidelleur X, Labanieh S, Lahitte P, Chauvel C (2010) The eruptive history of Morne Jacob volcano (Martinique Island, French West Indies): geochronology, geomorphology and geochemistry of the earliest volcanism in the recent Lesser Antilles arc. J Volcanol Geotherm Res 198:297–310 Giachetti T, Paris R, Kelfoun K, Ontowirjo B (2012) Tsunami hazard related to a flank collapse of Anak

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Computer Simulation of a Volcanic Debris Avalanche from Mt. Taranaki, New Zealand Jonathan N. Procter, Anke V. Zernack, and Shane J. Cronin

Abstract

More than 14 unconfined volcanic debris avalanche deposits are recorded in the last 210,000 years within the Mt. Taranaki ring plain demonstrating a high-magnitude but low-frequency hazard facing the surrounding community. The *7000 yr B.P. Opua Formation is the youngest of these events and exhibits the typical chaotic, polymodal, polylithologic and extremely poorly sorted characteristics of DADs. Despite the apparent invariance of these large-scale properties within the Opua deposit, the proportion of clay and sand dominated matrix to gravel/boulders clasts gradually changes from proximal to distal areas (>30 km from source) with the finer fractions, including clay contents, increasing with distance from the source. Scanning Electron Microscope analyses of 4 micron-grains show typical hackly textures and micro-cracks are common. There is no variation in the crack distribution or frequency in grains from different parts of the deposit, suggesting the cracking process occurred during the initiation phase of the debris avalanche. Analysis of the surface features of the deposit morphology shows

J. N. Procter (&)
A. V. Zernack
S. J. Cronin School of Agriculture and Environment, Massey University, Private Bag 11 222, Palmerston North, New Zealand e-mail: [email protected]

consistent variation with distance, in particular its mounds and hummocky surface. A near-source, initially chaotic surface gives way with distance to ridges of hummocks in flow-parallel direction. These eventually break-down into clusters of mounds and further to more widely spaced fields of individual mounds in distal areas. The Opua debris avalanche was generated by the gravitational collapse of a sector of the volcanic edifice that fragmented and flowed down a single catchment. Rapid changes in topography and slope resulted in the transformation of the flow into a more cohesive mobile body, which formed two major lobes marked by mound/hummock ridges. The granular flow model Titan2D was applied to evaluate possible emplacement conditions and collapse parameters. Titan2D, while useful for defining initial collapse parameters and major flow paths, could not adequately simulate the complex rheological transformations from a collapsing/sliding pile through a granular flow into a cohesive clay-rich flow with long runout and high apparent fluidity. It is also difficult to adequately define simulation parameters for this rapidly changing flow from the resulting geological deposits. Hence computer simulations of major flow paths must be used alongside insights from geological mapping to provide future-focussed hazard zones for debris avalanches.

© Springer Nature Switzerland AG 2021 M. Roverato et al. (eds.), Volcanic Debris Avalanches, Advances in Volcanology, https://doi.org/10.1007/978-3-030-57411-6_11

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Introduction

Catastrophic volcanic debris avalanches (VDA) present an infrequent yet highly destructive hazard on many stratovolcanoes worldwide (Siebert et al. 1987; Belousov et al. 1999). They dramatically alter the landscape, inundating surrounding volcanic ring plains and valleys with several tens of meters thick of volcanic debris (Procter et al. 2009). The infrequency of these events makes the observation of a volcanic debris avalanche extremely rare (Glicken 1996), necessitating a holistic approach to fully understand flow rheology and emplacement mechanisms of these events at any given volcanic centre. Prior to the 1980 eruption of Mt. St. Helens, volcanic diamictons emplaced by mass flow were commonly termed “lahars” without distinction between initiation and transport processes such as debris-flow versus flank collapse (e.g. Crandell 1971; Neall 1976a, b, 1979). The dramatic 18 May 1980 Mt. St. Helens demonstration of a volcanic debris avalanche and studies of the resulting deposits provided a detailed model for these volcano collapse events (Voight et al. 1981, 1983; Glicken 1998). Ongoing studies have differentiated the characteristics of debris avalanches and their deposits from other volcanic mass-flow types (e.g. Siebert 1984; Ui 1983, 1989; Crandell 1989; Glicken 1991, 1996). In addition, a series of models for volcanic mass wasting and volcanic ring-plain development have also been constructed (Palmer et al. 1991; Vallance et al. 1995; Vallance and Scott 1997; Belousov et al. 1999; van Wyk de Vries et al. 2000; Capra and Macias 2000, 2002; Capra et al. 2002; Waythomas and Wallace 2002; Shea et al. 2008). Debris avalanche processes have been inferred from interpreting features of their deposits such as shattered and jigsaw-cracked fragmental rock clasts, megaclasts, large rip-up clasts and matrix-properties as well as the shape of hummocks and mounds along the main axis of dispersal and in proximal areas (Palmer et al. 1991; van Wyk de Vries et al. 2000; Scott et al. 2001;

Shea et al. 2008). Scott et al. (2001) described debris avalanches as suddenly emplaced, very rapid flows of variably wet (but unsaturated) mixtures of rock fragments and soil formed in response to gravity. These are typically generated by large-volume landslides or slope failures (ranging from 0.05 to 45 km3) from the flanks of volcanoes (Varnes 1978; Ui 1983; Schuster and Crandell 1984; Pierson and Costa 1987). They may travel over 100 km from source and reach velocities as high as 100 ms−1 (Siebert et al. 1987; Crandell 1989; Scott et al. 2001). Debris avalanches are considered to behave initially as grain flows where particle collisions and frictional contact dominate during flow progression, in contrast to debris flows that can be supported by excess pore fluid pressures (Pierson and Costa 1987; Scott et al. 2001). Despite this, Scott et al. (2001) showed that debris avalanches can be complex and variable, ranging from pure dry grain flows in proximal reaches, transforming to large debris flows or rapid successions of these with distance. The general model of debris avalanche emplacement suggests an initial sliding motion of a solid or coherent rock mass and involves the non-turbulent transport of material with shearing occurring mainly along internal fractures (Siebe et al. 1992; Coussot and Meunier 1996; Aaron and Hungr 2016; Shea et al. 2008). The transformation to a flowing particulate mass usually takes place at the base of the volcano, or at a major break in slope following disaggregation of the material (Siebe et al. 1992). Further transformation into a cohesive debris flow occurs towards the final phases of transportation when water or water-saturated sediments and soils are incorporated (e.g. Palmer et al. 1991; Scott et al. 1995; Vallance and Scott 1997). The most common measure of debris avalanche runout estimations is based on the ratio of maximum height versus maximum runout (H/L ratio), which defines the coefficient of friction. This also provides the basis of many numerical models for simulation of flows (Ui et al. 2000). Models of rheology and emplacement have been developed to explain the relatively low dissipation and deformation of fragmental clasts that preserve original edifice stratigraphy (Takarada

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et al. 1999). These models incorporate features such as air/gas fluidized basal layers, mechanical fluidisation on lubricated substrates, acoustically or seismically triggered fluidization through to non-Newtonian, Bingham flow mechanisms (Legros 2002). Other models have focused on granular flow mechanisms involving low-density basal layers and low angles of (basal and internal) friction due to loss of mass through deposition or increased fluidization (Iverson et al. 1997). Ui et al. (2000) concluded that a combination of models probably applies to most volcanic debris avalanches depending on their composition, volume and environment.

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Computer Simulation of Volcanic Debris Avalanches

Modelling has become a common term to describe any physical property of a volcanic debris avalanche. In general models can be categorised into numerical, conceptual (Hungr 1995) and analogue (Andrade and van Wyk de Vries 2010; Paguican et al. 2012). These models can further be confused by the fact that they attempt to only characterise the initial collapse, the flowing mass, the emplacement of the mass or other sedimentological features. There is probably no single model or simulation that encompasses the entire spectrum of the phenomena and physical properties of a volcanic debris avalanche. In this study we focus on applying a computer simulation that attempts to simulate the collapsing mass and distribution of the resulting flowing mass which produces 3D outputs, providing data on the flow at each computational point. Numerical simulation and characterisation of mass flows is either based on determining empirical relationships identified from experiments and flow observations (Rickenmann 1995), or by comparing physics-based flow models to experimental or natural examples of flows (Denlinger and Iverson 2001). Computer models of flow hazards can be one to three dimensional, with the latter producing the most visually useful outputs for hazard analysis. All

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models attempt to simulate a flow of water (or any mass) from a discrete point to the next, or to reproduce several flow parameters across a complex natural terrain. The complexity of a numerical model increases as more physical characteristics are parameterised and included e.g. increasing the number of parameters that require constraining. Complexity and computational demand also strongly increase with the detail and resolution of the terrain representation over which the model runs, ranging from a simple series of cross-sections to a detailed submetre digital elevation model (DEM). The simplest of simulations assumed the granular flow to move as a single mass or sliding block, which implies flows have a constant viscosity and yield strength. To account for more complex frictional regimes, a Mohr–Coulomb model adaption was applied by Savage and Hutter (1989). Other simulations tend towards applying bulk continuum models, thus treating the entire phenomena is a single (fluid) phase (Denlinger 1987; Takahashi and Tsujimoto 2000). These models simulate grain flows separated by intergranular dispersive pressure in a frictional regime (Denlinger 1987). This was widely applied to represent a number of mass flows of low volume with travel distances unexplained by Mohr–Coulomb models. Recent continuum descriptions of flows in a depthaveraged, shallow-water framework have been applied by Denlinger and Iverson (2001) by adding in Coulomb frictional resistance. Simulations or models specifically focussed on volcanic debris avalanches are limited usually relying on applying or adapting numerical frameworks from the debris flow (Denlinger and Iverson 2001) or landslide community (Chen and Lee 2000; Chen et al. 2006). The work of Crosta et al. (2007, 2009) attempts to find a solution and simulation of the debris avalanche flow, entrainment and depositional processes by applying a combined Eulerian–Lagrangian method within a finite element modelling approach. Models such as that of Doyle et al. (2007) can account for two phases (granular and gaseous) and assumptions for the transfer of mass between them which may have some

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applicability to debris avalanches with material being exchanged between layers forming in the flow due to different frictional or shear regimes. Alternative developments are based on discrete element models of individual particular interaction which are also proving to be applicable to pyroclastic and mass flows (Bursik et al. 2005), provided large scale computing resources are available. The application of simulation tools such as Titan2D are becoming common e.g. Sheridan et al. (2005) and Murcia et al. (2010). New advances in simulating volcanic debris avalanches by applying discrete element modelling (Thompson et al. 2010) could bridge the gap between linking sedimentological observations, structural analysis, analogue modelling (Paguican et al. 2014) may provide more accurate simulations for hazard analysis. Recent simulation tools described by Iverson et al. (2016) utilising the D-Claw model provides greater accuracy in simulating phenomena of

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rapid shocks or runup and deflection of (debris) flows by the definition and incorporation of the Froude number, liquefication within the flow and the low basal friction. Given that a range of rheological conditions and flow mechanisms applicable to volcanic debris avalanches, modelling their future hazard potential by numerical simulation is very challenging. However, some of these hurdles must be overcome in order to provide the most realistic hazard assessment scenarios at frequently collapsing stratovolcanoes. To examine these constraints for modelling debris avalanches and debris flows a combined geological and modelling study was carried out at Mt. Taranaki using the most recent major debris avalanche unit (the c. 6700 yrs old Opua Formation; Neall 1979) as an example (Fig. 1). This single unconfined volcanic debris avalanche deposit displays a range of sedimentological variability with distance from source and the deposit margins and

Fig. 1 Location map of Mt. Taranaki and LandSat7 image of the ring-plain. Shown in yellow is the Opua Formation (Neall 1979). Indicated are also the location and identification numbers for grainsize samples and point count localities

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surface are well defined and observable. Using a combination of aerial photographic interpretation, GIS analysis and focussed sedimentological investigations, a model of emplacement for the Opua debris avalanche can be developed and compared to simulations based on the Titan2D granular flow model (cf. Patra et al. 2005). The simulation strategy, analysis of depositional features as modelling inputs applied in this case study are built upon the foundational work of Zernack (2009) and Procter (2009).

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Geological Setting of Mt. Taranaki/Egmont Volcano

The Taranaki volcanic succession consists of a group of Quaternary andesite volcanoes (Neall et al. 1986) that lie to the west of the major Taupo Volcanic Zone subduction-related volcanism (Cole 1986; Gamble et al. 1993; Wilson et al. 1995). The oldest of the Taranaki volcanic successions is Paritutu Volcano at the northwestern extreme, with a K‐Ar age of 1.75 Ma (Price et al. 1999). The next youngest volcano is c. 0.57 Ma old Kaitake, followed by Pouakai around 0.25 Ma (Neall 1979; Neall et al. 1986). Mt. Taranaki (Egmont) Volcano is the youngest and most southerly expression of the NW-SE trending volcanic alignment. Its activity began >210 ka (Zernack et al. 2011) with at least one eruption known to have occurred after CE1755 (Platz 2007). Mt. Taranaki and its ring plain are situated unconformably upon a marine basin sedimentary sequence of weakly consolidated sand-, silt- and mudstones with intercalated shell beds that record fluctuating sea levels from the Late Cretaceous to early Miocene (King and Thrasher 1996; Kamp et al. 2004). The sedimentary sequence is cut by numerous Quaternary faults, yet their common SSW-NNE orientation apparently reflects an earlier stress field (Sherburn and White 2006). An indicator of the more recent stress field may be given by the alignment of volcanic vents, because these tend to lie perpendicular to the extension direction (Nakamura 1977). Neall (1971) also identified a number of

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faults perpendicular to the current volcanic lineament particularly through Mt. Taranaki (the Oaonui Fault and Inglewood Fault). However, movements on these faults have not yet been correlated to any volcanic debris avalanche event. The current edifice of Mt. Taranaki rises to 2518 m above sea level (asl) and comprises c. 12 km3 (bulk volume) of lavas and pyroclastic deposits that are mostly younger than 14 ka (Neall 1979). The edifice is surrounded by an apron of volcaniclastic diamictons, deposits from debris avalanches, lahars, fluvial and pyroclastic deposits, that is an order of magnitude greater in volume than the Mt. Taranaki massif (>150 km3) (Neall et al. 1986; Zernack 2020—this volume). Volcanic debris avalanche deposits, attributed to the failure of former edifices, make up the largest component of volcaniclastic material within the near-circular Taranaki ring-plain. Four debris-avalanche deposits in west and south Taranaki were mapped by Neall (1979), including the Opua, Warea, Pungarehu and Stratford Formations and later Alloway et al. (2005) identified three additional debris avalanche deposits (Ngaere, Okawa and Motunui Formations) to the north-east and south-east of the volcano. Zernack et al. (2011) have revised this record to include 14 debris avalanche deposits over the entire record known for Mt. Taranaki. Palmer and Neall (1991) and Zernack et al. (2009, 2011) also recognised that these debris-avalanche deposits record large-scale destruction of the cone. In contrast, the intervals between their emplacement are characterised by phases of edifice (re)growth and periods of quiescence, represented by eruptive products and lahar deposits as well as paleosols and fluvial sediments, respectively (Zernack et al. 2009; Zernack 2020—this volume; Zernack and Procter 2020—this volume). Only a minimum runout distance can be estimated for the Taranaki debris avalanches because coastal erosion is continuously removing the toes of deposits around the Taranaki peninsula at c. 25–40 km from the edifice. These calculations indicate a minimum H/L ratio of 0.1. Uniquely, Mt. Taranaki volcanic debris

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avalanches could spread in an unconfined fashion onto the gently dissected ring plain, forming broad fans around the volcano (Neall 1979; Neall et al. 1986; Palmer and Neall 1991; Zernack et al. 2009). The largest known edifice failure at Mt. Taranaki produced the 20 ka Pungarehu Formation for which the onshore volume is estimated at 7.5 km3 (Ui et al. 1986) covering an area up to 250 km2. Its internal structure is characterised by two major components: fragmental rock clasts (FRCs) and matrix. FRCs are shattered or deformed pieces of lava and stratified volcaniclastic material that represent intact parts of the former volcanic edifice (Alloway et al. 2005). They are surrounded by interclast matrix, which includes clay-sized material. Mt. Taranaki lithologies, including those of FRCs within the avalanche deposits, range from vesicular red and black scoria through non-vesicular to holocrystalline, porphyric lavas of basaltic andesite to andesitic composition (Neall et al. 1986; Stewart et al. 1996). One of the smallest known debris avalanche units is the Opua Formation, estimated at only 0.35 km3 (Neall 1979) yet spreading over a large area and exhibiting similar structural features to the Pungarehu deposit. Mt. Taranaki has a high mean annual rainfall with a strong altitude-related gradient, ranging from 2400 mm on the lower flanks at c. 300 m asl., up to 8000 mm near the summit (Stewart et al. 1996). Palmer et al. (1991) highlights that most drainage originates from the upper slopes and mean discharge of all the rivers draining the volcano out to a 12 km radius equal 46 m3/s. This discharge is relatively constant and indicates a large amount of groundwater is located in the volcanic pile that may provide a considerable source of water during a collapse event as well as promoting weathering of the edifice interior. The incorporation of clays, in particular allophane and ferrihydrite (andic material), into inter-clast matrix is an additional significant factor in increasing the mobility of debris avalanches and related cohesive debris-flows (Alloway et al. 2005). Under the humid-temperate climatic conditions of the western North Island, abundant andesitic ash rapidly weathers to shortrange-order clay minerals such as allophane

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(Neall 1976a, b, c). Allophane consists of hollow spherules with diameters of 3.5–5 nm (Parfitt et al. 1981), and hence has a very high specific surface area and capacity for water retention (up to 300% of the weight of dry soil). During debris avalanche propagation, these cover-bed deposits will directly contribute allophane-rich, finegrained material into the flowing mass to greatly enhance its potential fluidity. Vallance and Scott (1997) suggested that clay-rich (cohesive) and wet flows might spread up to ten times farther than dry volcanic debris-avalanches.

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Debris Avalanche Nomenclature

Standard grain size terminology is applied in this study such as: clay (256 mm). Matrix refers to the components of the deposit that make up the finest grainsize (4 mm). Clay refers to clay-sized particles as well as crystalline and short-range-order clay minerals. Debris avalanche deposits are distinguished from other volcaniclastic diamictons with similar textural characteristics by their content of mega-clasts (Neall 1979; Palmer et al. 1991; Scott et al. 2001), which are single (encased) definable fragments of lithologically or stratigraphically coherent material derived from the original edifice. Other diagnostic features include jig-saw fractured clasts, major zones of strongly sheared and deformed soft clasts and sheared basal margins with common rip-up clasts (Siebert et al. 1987; Glicken 1991; Ui et al. 2000; Scott et al. 2001). The internal structure of Taranaki debris avalanche deposits is characterised by fragmental rock clasts (FRC) and matrix. The matrix is unsorted and unstratified and may contain rip-up clasts of plastically distorted soil, peat and tephra layers as well as wood fragments derived from the terrain beneath. Intra-clast matrix is a separate entity that occurs within FRCs and

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megaclasts of original breccias from the edifice (Alloway et al. 2005). A FRC is defined as a fragmented or deformed piece of lava or layered volcaniclastic material commonly preserving stratification and/or contacts formed within the original volcanic edifice (Alloway et al. 2005). Scott et al. (2001) and Palmer et al. (1991) define megaclasts as being >1 m in diameter. These can be large coherent fragments of rock, intact portions of the original edifice strata, fragmented clasts that have been partially disaggregated and those that have completely disaggregated to form a “domain” of rock fragments that are recognisably related and could represent a jigsaw puzzle (Shreve 1968; Gaylord et al. 1993). The lithology and ratio of matrix to FRCs can vary within each debris-avalanche deposit depending on the characteristics of the source area, related volcanic activity, as well as flow rheology and interaction with paleo-physiography. The content of clay and/or clay-sized particles is also a diagnostic feature of debris avalanche deposits and can provide information on source areas, initiation and emplacement processes. Wet debris avalanches and clay-rich (cohesive) debris flows are typically associated with the collapse of fluid-saturated portions of a volcanic edifice (Vallance and Scott 1997). The available water, pore-water and alteration or weakness of the preavalanche mass contributes to the rapid transformation from debris avalanche to clay-rich debris flow.

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The Opua Formation

The Opua Formation (Fig. 1) was first identified by Neall (1979) as a debris avalanche deposit on the south western ringplain that consists of a poorly sorted, coarse, brecciated axial facies flanked by marginal conglomerate and sands. The deposit forms a broad fan (7 km wide) of c. 4 m thickness of deposit resulting in a volume of 0.35 km3, a similar volume to that of the amphitheatre of the source area on the southern volcano flanks. The deposit exhibits a characteristic hummocky landscape along its axial facies with a higher density of larger hills along

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the main dispersal axis that reduce in spatial density and size laterally and longitudinally. Neall (1979) also considered the largest mounds to be protruding from beneath the Opua Formation and represented mounds from earlier deposits. The marginal facies does not exhibit mounds and was thought to have been emplaced by debris flows. 14C dates of wood from within the deposit returned an age of 6570 ± 110 yrs B. P. (Neall 1979). Three facies were recognised in the deposit; axial a, axial b, and marginal facies (Neall 1979; Alloway 1989; Palmer et al. 1991; Alloway et al. 2005). • Axial a facies is where fragmental rock clasts dominate and interclast matrix occupies 90 vol% of the deposit, and no mounds or hills occur. In exposure, the Opua debris avalanche deposit is a poorly-sorted diamicton consisting of a characteristic yellow to brown, poorly sorted, silty or clay-rich sand matrix with a variety of andesitic to basaltic-andesite fragmental rock clasts (Fig. 2). The internal stratigraphy generally remains consistent with distance from the source. In distal areas the matrix exhibits higher contents of pumice and/or a wet, greasy clay-rich texture. The matrix is mainly composed of rock fragments of the same composition of the clasts, along with minor pumice fragments and pyroxene, hornblende and plagioclase crystals. The clasts are angular to sub-rounded, porphyritic, basalticandesite to andesite lithologies, ranging in colour from black and grey to red. Pumice clasts are

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typically weathered and sub-rounded. They occur in low proportions and are always