Iceland Within the Northern Atlantic, Volume 2: Interactions between Volcanoes and Glaciers [1 ed.] 1789450152, 9781789450156

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Iceland Within the Northern Atlantic 2

SCIENCES Geoscience, Field Director – Yves Lagabrielle Lithosphere-Asthenosphere Interactions, Subject Head – René Maury

Iceland Within the Northern Atlantic 2 Interactions between Volcanoes and Glaciers

Coordinated by

Brigitte Van Vliet-Lanoë

First published 2021 in Great Britain and the United States by ISTE Ltd and John Wiley & Sons, Inc.

Apart from any fair dealing for the purposes of research or private study, or criticism or review, as permitted under the Copyright, Designs and Patents Act 1988, this publication may only be reproduced, stored or transmitted, in any form or by any means, with the prior permission in writing of the publishers, or in the case of reprographic reproduction in accordance with the terms and licenses issued by the CLA. Enquiries concerning reproduction outside these terms should be sent to the publishers at the undermentioned address: ISTE Ltd 27-37 St George’s Road London SW19 4EU UK

John Wiley & Sons, Inc. 111 River Street Hoboken, NJ 07030 USA

www.iste.co.uk

www.wiley.com

© ISTE Ltd 2021 The rights of Brigitte Van Vliet-Lanoë to be identified as the author of this work have been asserted by her in accordance with the Copyright, Designs and Patents Act 1988. Library of Congress Control Number: 2021934743 British Library Cataloguing-in-Publication Data A CIP record for this book is available from the British Library ISBN 978-1-78945-015-6 ERC code: PE10 Earth System Science PE10_5 Geology, tectonics, volcanology PE10_13 Physical geography PE10_18 Cryosphere, dynamics of snow and ice cover, sea ice, permafrosts and ice sheets

Contents List of Abbreviations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brigitte VAN VLIET-LANOË and Françoise BERGERAT

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Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brigitte VAN VLIET-LANOË and René MAURY

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Chapter 1. Young Icelandic Volcanism and its Implications . . . . . René MAURY and Brigitte VAN VLIET-LANOË 1.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2. Icelandic magma series . . . . . . . . . . . . . . . . . . . . . . 1.2.1. Lava types . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2.2. Geochemical diversity of young Icelandic basalts and their sources . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2.3. Some geochemical constraints concerning the origin and geodynamic evolution of Iceland . . . . . . . . . . . . . . 1.3. Central volcanoes and active fissural systems . . . . . . . . 1.3.1. Central volcanoes . . . . . . . . . . . . . . . . . . . . . . 1.3.2. Fissural volcanism and subaerial lava flows . . . . . . 1.3.3. Hydromagmatism . . . . . . . . . . . . . . . . . . . . . . 1.4. Volcanic hazards in Iceland . . . . . . . . . . . . . . . . . . . 1.4.1. Hazards related to lava flows . . . . . . . . . . . . . . .

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1.4.2. Hazards related to explosions and gas emissions 1.4.3. Jökulhlaups and associated hazards . . . . . . . . 1.4.4. Icelandic dust: a consequence of volcanism . . . 1.5. References . . . . . . . . . . . . . . . . . . . . . . . . . .

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Chapter 2. Volcanism and Glaciations: Forcings and Chronometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hervé GUILLOU, René MAURY and Brigitte VAN VLIET-LANOË 2.1. Subglacial volcanic landforms . . . . . . . . . . . . . . . . . . . 2.1.1. Subglacial isolated volcanoes or tuyas . . . . . . . . . . 2.1.2. Hyaloclastite ridges or tindar . . . . . . . . . . . . . . . . 2.2. Volcanism, deglaciation and climate . . . . . . . . . . . . . . . 2.2.1. General features: deglaciation, discharge and partial melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2.2. Deglaciation and climate feedback . . . . . . . . . . . . . 2.3. The hypothesis of a link between volcanism and climate and its test by dating . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.1. The K-Ar chronometer . . . . . . . . . . . . . . . . . . . . 2.3.2. The combination of K-Ar and 40Ar/39Ar methods for dating Icelandic volcanism . . . . . . . . . . . . . . . . . . . 2.3.3. A link between volcanism and climate according to K-Ar ages? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3.4. A rhyolitic volcanism synchronous with deglaciations? 2.4. References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Chapter 3. Cenozoic Evolution of Iceland and the Cryosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brigitte VAN VLIET-LANOË and Hervé GUILLOU 3.1. Ice ages and the opening of the Atlantic . . . . . . . . . 3.1.1. The Middle and Final Miocene cooling . . . . . . 3.1.2. The acceleration of the Middle Pliocene . . . . . 3.1.3. The Middle Pleistocene Transition . . . . . . . . . 3.1.4. The initiation of thermohaline circulation. . . . . 3.2. Iceland’s Quaternary glaciations . . . . . . . . . . . . . 3.2.1. Conditions for the development and functioning of ice caps . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2.2. Glacio-isostasy . . . . . . . . . . . . . . . . . . . . 3.2.3. Icelandic data . . . . . . . . . . . . . . . . . . . . . 3.2.4. The Icelandic record . . . . . . . . . . . . . . . . .

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Contents

3.3. The last glacial episode and its deglaciation 3.3.1. The Weichselian. . . . . . . . . . . . . . 3.3.2. The Last Glacial Maximum . . . . . . . 3.3.3. Deglaciation and the Holocene . . . . . 3.4. Iceland today, its climate and vegetation . . 3.4.1. The climate . . . . . . . . . . . . . . . . . 3.4.2. Ocean circulation and climate. . . . . . 3.4.3. Soil, people and climate . . . . . . . . . 3.4.4. Soils and erosion . . . . . . . . . . . . . 3.5. References . . . . . . . . . . . . . . . . . . . .

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142 142 144 150 161 161 162 167 173 179

Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brigitte VAN VLIET-LANOË, René MAURY and Hervé GUILLOU

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References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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List of Authors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

215

Summary of Volume 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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List of Abbreviations σ1, σ2 and σ3

Maximum, intermediate and minimum principal stresses of the stress tensor

σHmax

Maximum horizontal stress

A AMO

Atlantic multidecadal oscillation

B BTVP

British Tertiary Volcanic Province

C CGFZ

Charlie–Gibbs Fracture Zone

CGPS

Communicative Global Positioning System

D DL

Dalvik Line

DMM

Depleted MORB mantle

DO

Dansgaard–Oeschger event Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021.

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DSOW

Denmark Strait overflow water

DTM

Digital terrain model

E E-MORB

Enriched mid-ocean ridge basalts

EM

Enriched mantle

EUR

Europe

EVZ

East Volcanic Zone

F FLF

Flat-lying flows

G GEBCO

General Bathymetric Chart of the Oceans

GIA

Glacio-isostatic adjustment

GIFR

Greenland–Iceland–Faroe Ridge

GIR

Greenland–Iceland Ridge

GL

Grimsey Line

GPS

Global Positioning System

H HFF

Húsavík-Flatey Fault

HIMU

High Mu mantle (Mu = U/Pb)

List of Abbreviations

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I ICPMS

Inductively coupled plasma mass spectrometry

IFR

Iceland–Faroe Ridge

IGS

International GPS Service

IMO

Icelandic Meteorological Office (Veðurstofa Íslands)

InSAR

Interferometric Synthetic Aperture Radar

IRD

Ice-rafted detritus

ISOW

Iceland–Scotland Overflow Water

ÍSNET

GPS Network surveys of the National Land Survey of Iceland (Landmælingar Íslands)

J JMFZ

Jan Mayen Fracture Zone

K KR

Kolbeinsey Ridge

L LBA

Labrador–Baffin axis

LGM

Last Glacial Maximum (extension)

LIP

Large igneous provinces

M M or MW

Moment magnitude

MAR

Mid-Atlantic Ridge

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Mb

Body-wave magnitude

ML

Local magnitude

MS

Surface-wave magnitude

N N-MORB

Normal mid-ocean ridge basalts (depleted)

NADW

North Atlantic Deep Water

NAIP

North Atlantic Igneous Province

NAM

North America

NEIC

National Earthquake Information Center (United States)

NGRIP

North Greenland Ice Core Project

NVZ

North Volcanic Zone

O OIB

Ocean island basalts

OSC

Overlapping spreading center

R RP

Reykjanes Peninsula

RR

Reykjanes Ridge

S SDRs

Seaward-dipping reflectors

SIL

South Iceland Lowland network

SISZ

South Iceland Seismic Zone

List of Abbreviations

T TFZ

Tjörnes Fracture Zone

U USGS

United States Geological Survey

W WVZ

West Volcanic Zone

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Preface Brigitte VAN VLIET-LANOË and Françoise BERGERAT

This collective work is the logical conclusion of more than 30 years of French research in Iceland, with the support of various programs and institutions. It has also benefitted from the contribution of a CNRS Thematic School on Iceland, which was held in Brest in 2010 and which was strongly impacted by the eruption of Eyjafjallajökull. This book is the fruit of the work of a group of complementary researchers who are very fond of Iceland. Our thoughts turn to Jacques Angelier who left this basaltic ship a little too early. There are multiple authors to each chapter – with a principal author for each one – in order to provide a multidisciplinary approach to the discussed scientific problems and take into account all our publications up to the most recent ones (2019–2020). French research in Iceland began in the mid-1980s, initiated by Françoise Bergerat (Sorbonne Université, formerly Université Pierre et Marie Curie, in Paris) in search of an “emerging oceanic ridge”, in collaboration with Jacques Angelier†, then Catherine Homberg. Very quickly, this collaboration was extended to Icelandic colleagues, Águst Guðmundsson (London), Kristjan Sæmundsson, Ragnar Stefánsson and Sigurdur Rögnvaldsson†. The first work focused on the analysis of brittle deformations and then turned to sismotectonics.

Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021.

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This work was then supplemented, from the 2000s, by the geodetic campaigns of the team from the Université de Savoie in Chambéry led by Thierry Villemin in collaboration with Halldór Geirsson and his group. At the beginning of the 1990s, Laurent Geoffroy began (in Paris) work on the Thule basaltic provinces (Scotland, Ireland, Faroe Islands), continued from the 2000s (at the Université du Maine, in Le Mans) on the other side of the Atlantic, in Greenland. The analysis of the morphology of Iceland began in the mid-1990s at the Université de Rennes-I, with Olivier Dauteuil and Brigitte Van Vliet-Lanoë, and then extended to the neighboring ocean in relation to volcanism and the evolution of the North Atlantic. At the same time, the Neogene and Quaternary climatic history of the island, recorded by stratigraphy, was consolidated with dating carried out by Hervé Guillou and his colleagues and by geochemistry carried out at the Université de Bretagne Occidentale, in Brest, in close collaboration with Águst Guðmundsson (Hafnafjördur), Kristjan Sæmundsson and Helgi Björnsson’s team. The last stage of this work is currently being developed in the Géosciences Océan laboratory in Brest, with Laurent Geoffroy and René Maury. It concerns the evolution of the North Atlantic based on Icelandic and Greenlandic data. The material and logistical support of the Icelandic authorities proved to be very constructive both for field work and for data acquisition and sharing: IMO (Veðurstofa Íslands/Icelandic Meteorological Office); ISOR (Íslenskar orkurannsóknir/Icelandic energy research), formerly Orkustofnun (National Energy Authority); Landsvirkjun (National Power Company) and Vatnajökull National Park. This research would not have been as fruitful without the physical and intellectual help of all our students, at Master’s level and/or with their thesis works: Olivier Bourgeois, Magalie Bellou, Jean-Christophe Embry, Loïc Fourel, Sebastian Garcia, Guillaume Gosselin, Solène Guégan, Romain Plateaux, Lionel Sonnette, Anne Sophie Van Cauwenberge, Ségolène Verrier and Audrey Wayolle. Finally, this work was made possible because of the assistance of the French Embassy in Iceland and funding from the European Commission, 4th and 5th PCRD (PRENLAB-1 and -2, PREPARED and SMSITES programs); the Paul Émile Victor Institute (IPEV), formerly the Institut français pour la recherche et la technologie

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polaires (IFRTP) (Arctic Program 316); the Icelandic Ministry of Education; and the French Ministry of Foreign Affairs (Franco-Icelandic scientific and cultural collaboration program). We also thank Bernadette Coleno, Marion Jaud, Laurent Gernigon and Alexandre Lethiers for their contributions to the figures in this volume. March 2021

Introduction Brigitte VAN VLIET-LANOË and René MAURY

Figure I.1. Iceland from Space (document Geographical Institute of Iceland/Landmælingar Ísland [LMIs])

For color versions of the figures in this Introduction see, www.iste.co.uk/vanvliet/iceland1.zip. Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021.

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Iceland (Figure I.1), a young and isolated island in the middle of the Atlantic Ocean, has only very recently been discovered in terms of the scale of human history. Irish monks (the Papar) passed from island to island in their curraghs (Figure I.2) via the Shetland Islands and the Faroe Islands to evangelize the legendary Hyperborea. These journeys took place as early as the 6th century, a period with cold volcanic winters. The Papar discovered a world of fire and ice, the gates of hell. They settled in round peat-covered huts and dug shovel caves in the consolidated sandy interglacial formations in the south of the island.

Figure I.2. (A) Icelandic stamp illustrating the discovery of Iceland by Irish monks (Papar). (B) The glacial lake Jökulsarlón dominated by the volcano Oræfajökull (Brigitte Van Vliet-Lanoë©)

Two hundred years later, the Vikings, warriors but also more than anything farmers in search of cultivable land (Figure I.3), settled in the south and west of the island from 860 AD (for Anno Domini, year 1 of the Christian calendar) on wooded land made fertile by thick layers of volcanic loess. This is the landmana of the Icelandic sagas. They installed their parliament, the Alþing, around 900 AD, in a remarkable site (Figure I.4), which became a high place of plate tectonics, the Þingvellir graben, the boundary between the European and American plates. These fertile lands were surmounted for at least 400,000 years by a fire monster, the Hekla volcano (Figure I.5). Its Plinian eruption in 1104 AD (H1, volcanic explosivity index of 5) destroyed many Viking settlements in the Rangavellir, not only by falling pumice and gas but also by the associated glacial megafloods, the jökulhlaups, submerging the Þjorsárdalur with a wave of muddy water more than 25 m high. At that time, the Hekla must have been more ice-covered than it is today.

Introduction

Figure I.3. Traditional sheepfolds in southern Iceland with the volcano Hekla in the background (Rangavellir) (Brigitte Van Vliet-Lanoë©)

Figure I.4. (A) The Þingvellir graben, seen from an airplane (Þingvellir National Park Web site). The Alþing site is located in front of the white buildings (chapel). (B) Reconstruction of the Alþing in the Middle Ages by W.G. Collingwood (1897)

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Figure I.5. (A) The Hekla volcano (Brigitte Van Vliet-Lanoë©) and (B) its cartographic representation on the map Islandia of Ortelius (1585)

Despite the island’s long isolation from continental Europe, there is a lot of information about its history. Indeed, Icelanders have jealously preserved their language and ancient books, including the famous sagas, and have often proved to be great writers and avid readers, even on isolated farms.

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In addition to their literary and historical interest, the sagas represent a source of exceptional paleo-environmental information on a period whose climatic evolution was very complex: the Medieval Optimum and the climatic degradation that followed. The University of Iceland was founded in 1911 and, due to its special nature, Iceland is the country with the highest proportion of geologists and especially volcanologists among its population.

Figure I.6. Typical landscape of ancient basalts on the eastern coast of Iceland (Skridalur) (Brigitte Van Vliet-Lanoë©)

Iceland is a land of fire and ice, still sparsely populated (about 350,000 inhabitants in 2020), prized by tourists for its “unspoiled”, photogenic character and its many natural wonders, although Viking colonization quickly made the forest disappear. But recent tourist development has also caused an invasion of 4×4 vehicles, brand new hotels and vacation huts, raising the standard of living of the population, but gradually destroying a natural heritage – including the geological heritage – surprisingly well preserved until the early 21st century. Industrial development (geothermal, hydroelectricity and electrometallurgy) kept the population in the peripheral sectors of the island and above all modified the landscape of the coastal zones. Whatever one does or looks at in Iceland is de facto connected to the geological history of the island (Figure I.6). Despite its remoteness, Iceland is a land that directly influences Western Europe through its position in the north-central Atlantic, as a beacon of the Gulf Stream and thermohaline circulation, or through its meteorological depression. But it is also a

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land consisting mainly of layered basaltic piles, still active from a tectonic and volcanic point of view. We were reminded of it by the last eruption of the Eyjafjallajökull (March–October 2010) with its plume of ash that invaded Europe and disrupted intercontinental commercial flights (Figures I.7 and I.8).

Figure I.7. The eruption of the Eyjafjallajökull in 2010: its jökulhlaup (Reykjavik Helicopter©) and plume (Earth Observatory, NASA)

Figure I.8. Sheep disturbed by the ash from the eruption (Flickr©)

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The glaciers are located on volcanic edifices, considered to be at least Quaternary. The largest ice cap, the Vatnajökull, rests on some of the most active volcanoes of the island, located above the summit of a deep magma plume. Bárðarbunga (Figure I.9) is one of the volcanoes found above the Icelandic hotspot and is located on the western margin of the present Vatnajökull ice cap.

Figure I.9. Digital terrain model of Vatnajökull (black line: current cap boundary) completed with the flood drainage positioning and the potential extension (to a depth of 200 km) of the Icelandic mantle plume (in gray) (source: H. Björnsson, 2009)

The most recent eruption of this volcano (August 2014–February 2015; Figure I.10) was linked to the draining of a magma chamber located 12 km below the caldera, following the climate driven melting of the cap (about 1 m/year). To the northwest, the most impressive lava flow since the 18th century, the Holhurhaun flow, occurred along a fracture line, in association with swarms of earthquakes that stretched to the Askja volcano in the north. The previous eruption, that of Veiðivötn, had flown toward the south in 1747, awakening the Torfa volcano at the same time.

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The Bárðarbunga is also a source for jökulhlaups or megafloods, resulting from the melting of the glaciers by the heat of the lava emitted and which mostly flow toward the north.

Figure I.10. (A and B) Views of the Bárðarbunga caldera obliterated by ice during the 2014 eruption with a melting cauldron to the west (A and black arrow) (photo: mbl.is/RAX). (C) Satellite image of the emersion of the Holhurhaun flow (star) on August 13, 2014 at the foot of the Bárðarbunga (B) (photo: TerraSAR-X)

Another major volcanic structure is located in the center of the ice cap, directly above the top of the mantle plume: it is the triple caldera of Grímsvötn (Figures I.9 and I.11), which emitted the vast majority of basaltic tephra that hide the glaciers and reach the lands surrounding the North Atlantic.

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The most famous is the Saksunarvatn tephra splayed around 10,200 years cal BP. This volcano is never at rest; its current eruptive frequency is about 10 years and it also remained continuously active during the Ice Age, but with a lower frequency. It is mainly responsible for the formation of subglacial lakes and is at the origin of most of the jökulhlaups that gully the emissaries of the Vatnajökull cap (Figures I.12 and I.13). At present, these floods mainly destroy road infrastructures such as the Main Highway (N1).

Figure I.11. Rim of the northern caldera of Grìmsvötn (Ragnar Sigurdsson©)

In northern Iceland, volcanic activity is also significant, in association with the northern rift. Many geothermal fields are exploited there, such as the Krafla field northeast of Lake Myvatn (Figures I.14 and I.15). This volcanic activity also occurs at sea, both in the north in the Kolbeinsey Ridge and its intermittent island (white point in Figure I.16(A)) and in the southwest along the Reykjanes Ridge, or in the Vestmann Islands, a southern extension of the East Volcanic Zone.

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Figure I.12. The Grìmsvötn volcano. (A) Initiation of the northward flow associated with a collapse of the ice mass (sun to the west), which led to the great jökulhlaup of November 1996 (Oddur Sigurðsson©). (B) Grimsvötn crater at the end of the 2011 eruption (Dima Moiseenko©). (C) Interstratified and deformed basaltic tephras in the terminal glacier tongue of Brúarjökull (LMIs)

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Figure I.13. (A) The jökulhlaups of the Skafta River from the Grimsvötn in 1996 (M.T. Gudmundsson©) and (B) Main Highway (N1) in 2011 (Veðurstofa Íslands©), frequently repaired since 1970, with (C) the jökulhlaup memorial of November 1996: two enormous pieces of the metal deck of the old bridge, twisted like common wires (Françoise Bergerat©)

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Figure I.14. (A) Fissural eruption of Krafla in 1980, along fractures arranged en échelon. (B and C) The geothermal power plant (C) narrowly escaped destruction by lava flows (flow with white arrow) (B-C: Brigitte Van Vliet-Lanoë©)

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Figure I.15. Eruption of Krafla in 1980: hornitos on fractures and lava flows in 1997 (Brigitte Van Vliet-Lanoë©)

The latter were the locus of a first submarine eruption in 1963 (building of the Surtsey volcano), then of a fissural eruption (followed by a strombolian phase) partially destroying the town of Vestmanayer on the main island of Heimaey in 1973. In relation to volcanic activity and especially to tectonic activity, seismicity is permanent in Iceland and major earthquakes have regularly occurred, particularly in the northern peninsulas (Húsavík region) and in the whole south of the island. When crossing the lava fields between Hveragerði and the Hekla, many remarkably preserved traces of major historical earthquakes (M > 6) can be observed (Figure I.17).

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Figure I.16. (A) The submarine ridges of Kolbeinsey and (B) of Reykjanes (multibeam echosounder images, HAFRO.is). (C) The eruption of Heimaey, building the “Mountain of Fire” (Eldfell) in 1973

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Figure I.17. Trace of the Réttarnes seismic fault (1294 or 1732) in the Rangavellir: South Iceland Seismic Zone (Françoise Bergerat©)

If in the north of the island the current earthquakes occur mainly offshore, the Húsavík and Kopasker agglomerations are however far from being sheltered from a significant seismic event, and in the south, several major earthquakes have occurred very recently (Mw 6–7; June 2000, May 2008). While Icelandic houses are relatively insensitive to earthquakes (Figure I.18), the same cannot be said for road infrastructure or greenhouses. The temporary rise or fall of water tables or lakes is frequent, reactivating or deactivating geysers and causing fluid escapes. This is particularly the case in the Hveragerði or Geysir region: Strokkur is currently the most active and Great Geysir is currently intermittent (Figure I.19). The cold pole of Iceland is represented by its glaciers, currently relatively little extended but which covered practically all the island at the time of the last glaciation, inhibiting the activity of a great majority of the volcanoes. Most of the time, they settled at the top of the volcanic edifices constituting the high points of the island such as Vatnajökull (2,009 m at Bárðarbunga) or Hofsjökull (1,765 m at Habunga; Figure I.20).

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Figure I.18. Destruction caused during the earthquakes of June 2000 in Bitra: (A and B) farm buildings, (C) main highway (N1) (A-B-C Françoise Bergerat©), and at the end of May 2008 in Hveragerði: (D) dislocated pipes and damaged greenhouses (Brigitte Van Vliet-Lanoë©)

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Figure I.19. Successive phases of an explosion of the Strokkur geyser, Geysir geothermal field (Brigitte Van Vliet-Lanoë©)

Figure I.20. The Hofsjökull. Document made from radar images (CNES©). The caldera is located at the top left of the picture

These glaciers have profoundly carved the island since the Neogene, with deep glacial valleys, ice-smoothed or striated rocks, countless drumlins and large areas of abandoned glacial sediments on the central plateau, especially around the Kerlingarfjöll (Figure I.21). Some volcanoes have typically subglacial morphologies, such as tabular volcanoes or tuyas, the best known of which is Herðubreið (Figure I.22). Others form alignments of ridges, the tindar, which formed at the margin of the melting caps (Figure I.23).

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The waters from these glaciers have also shaped canyons with huge waterfalls, on powerful, gray and loaded water rivers, the jökullsá (Figures I.23 and I.24). These waters are currently collected for an important hydroelectric production with mainly industrial purposes (aluminum and rare metals extracted from imported ores). This resource accounts for 72% of Iceland’s electricity production. Various cap outlets are currently being developed and managed, with water stored in very large dams, generally superimposed on the same course and designed to resist jökulhlaups of interglacial rank. Global warming in recent decades and potentially induced volcanism are likely to call this policy into question.

Figure I.21. (A) The Kerlingarfjöll surrounded by its glacial desert. (B) Perched upper cirque and (C) ice-smoothed rocks of the eastern fjords (Mjóifjörður, south of Seiðifjörður) (Brigitte Van Vliet-Lanoë©)

Introduction

Figure I.22. A subglacial tabular volcano: the Herðubreið, north of Vatnajökull, North Volcanic Zone (Brigitte Van Vliet-Lanoë©)

Figure I.23. Jökulsá á Kreppa north of Vatnajökull with hyaloclastite or tindar ridges (Brigitte Van Vliet-Lanoë©)

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Figure I.24. A key Icelandic resource: water. (A) Bruarjökull outlet (the glacier is at the bottom of the photograph) (LMIs©). (B) One of the Dettifoss waterfalls (Jökulsá á Fjöllum). (C) The Haslsón dam on the Jökulsá á Brú. (D) The Fannahlið aluminum plant (Hvalfjörður) (photos B, C and D: Brigitte Van Vliet-Lanoë©)

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On land, Iceland’s only important and renewable resources are its water and, as a result, its hydroelectricity, as well as its many geothermal sites related to the presence of the hot spot. In this two-volume book, we will present the geological and glacial history of this island, its current tectonic and volcanic activity and the impact of its formation on the climatic evolution of the last few millions of years. Volume 1 replaces Iceland within the geological framework of the North Atlantic, and describes its tectonic and geodynamic evolution. This second volume is dedicated to the study of the interactions between Icelandic volcanism and external geodynamics, i.e. with glaciations and the climatic evolution of the Atlantic zone during the Neogene and the Quaternary.

1

Young Icelandic Volcanism and its Implications René MAURY and Brigitte VAN VLIET-LANOË

1.1. Introduction Iceland is an atypical volcanic island resulting from the superposition of the Mid-Atlantic oceanic ridge and a hotspot. Volcanism was however disturbed at least three million years ago by the onset of glaciations (Chapters 2 and 3) which provoked the formation of large ice caps. The island has a normal volcanic functioning during interglacial periods. While this volcanism is partially inhibited at the maximum extent of the glaciers because of the loading exerted by the ice caps (glacio-isostasy), it is on the other hand strongly disturbed during periods of unloading related to a partial or total deglaciation of the island. This last phenomenon influences the volume and the process of lava emission as well as the type of volcanism, sub-glacial or hydromagmatic, and the geochemical evolution of magmas. It will be discussed in Chapter 2. The assessment of Icelandic magma production since its emergence 16 My ago is problematic because the island is part of the Greenland–Iceland–Faroe Ridge (Figure 1.5 in Volume 1), which has an average crustal thickness of 30 km (section 3.3.1 and Figure 3.20 in Volume 1) that reaches 40 km below Iceland (Figure 3.23(a) in Volume 1).

For color version of the figures in this chapter see www.iste.co.uk/vanvliet/iceland2.zip Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021. Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

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Recent studies tend to interpret the lower part of this crust as a stretched continental crust (section 3.3.2 of Volume 1) perhaps corresponding to an ancient Caledonian suture (Figure 3.9 of Volume 1). Only its upper seismogenic part, 7 km thick (Foulger et al. 2003), is most probably of pure volcanic origin.

Figure 1.1. Distribution of recent Icelandic volcanism (modified from Þórðarson and Höskuldsson 2002)

COMMENT ON FIGURE 1.1. – Main central volcanoes and active fissure systems. 1: Reykjanes; 2: Krýsuvík; 3: Brennisteinfjöll; 4: Hengill; 5: Hrómundartindur; 6: Grímsnes; 7: Hrafnabjörg; 8: Prestahnjúkur; 9: Kjölur; 10: Hofsjökull; 11: Kerlingarfjöll; 12: Tungnafellsjökull; 13: Vestmannaeyjar; 14: Eyjafjallajökull; 15: Katla; 16: Tindfjöll; 17: Hekla; 18: Torfajökull; 19: Veiðivötn; 20: Grímsvötn; 21: Kverkfjöll; 22: Askja; 23: Fremrinamur; 24: Krafla; 25: Þeistareykir; 26: Öræfajökull; 27: Esjufjöll; 28: Snæfell; 29: Ljósufjöll; 30: Lýsuskarð;

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31: Snæfellsjökull. Magmatic affinities (tholeiitic, transitional and alkaline) according to Sigmarsson and Steinthorsson (2007) and Sigmarsson et al. (2008). If we consider that the presently emerged part, Neogene in age, constitutes the summit of the island, with an outlined area of 120,000 km2 (Figure 1.2 of Volume 1), we obtain a Neogene and Quaternary magmatic volume of 840,000 km3. Adding to this the volume of deep intrusions, very difficult to estimate but often considered to be about half of that of the products emitted at the surface in an extensive context, the total volume (1.26.106 km3) of magma emitted/injected during the last 16 My corresponds to a production rate of about 80,000 km3/My. The latter is lower than that of the Hawaiian hotspot during the last million years (Big Island construction, 213,000 km3), but much higher than the estimated rates for those of Réunion and Tahiti (about 20,000 km3/My), and even higher than those of the Canary Islands and the various Polynesian archipelagos (Chauvel et al. 2012). During the Holocene, after deglaciation, the average volume of lava emitted at the surface, globally lower (20 km3/1,000 years) (Hjartason 2003) has nevertheless remained equivalent to that of the Réunion hotspot. Icelandic magma reservoirs are the result of the coalescence of fault swarms (Guðmundsson 2000) either shallow (about 2–3 km deep, e.g. shallow Grímsvötn; Krafla) or deeper and in basicrustal or even infracrustal position in some cases (about 12–20 km, at the fragile-ductile limit of the lithosphere: Hekla, Bárðarbunga and deep Grímsvötn). They are mainly located within the rifts connected with the Mid-Atlantic Ridge, such as that of the Reykjanes Peninsula (RP), and are fed laterally from the Icelandic hotspot. They give rise to large fissural volcanic fields and central volcanoes, mostly located at the level of the rifts and the transforming zones that connect them (Figure 4.1), although “off-axis” volcanism is not uncommon (Snæffels). The chemical composition of Icelandic basaltic magmas differs from that of the depleted oceanic basalts (MORB-N) of the Mid-Atlantic Ridge by the presence of enrichments coming from mantle components of the plume. Moreover, it is locally disturbed by contamination due either to a previously emplaced and altered oceanic basaltic crust, or to slivers of continental crust, also altered, remaining at the base of the Icelandic plateau (Chapter 1 of Volume 1). 1.2. Icelandic magma series 1.2.1. Lava types Recent lavas in Iceland are mainly basalts, of which three main geochemical types can be distinguished: tholeiitic, alkaline basalts and transitional basalts (Figures 1.1

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and 1.2). Tholeiitic basalts or tholeiites, by far the more frequent, are also geochemically closest to the depleted oceanic basalts (MORB-N) of the Mid-Atlantic Ridge, and all transitions between the two types are observable along the Reykjanes and Kolbeinsey Ridges (RR and KR) when approaching Iceland (Agranier et al. 2005; Blichert-Toft et al. 2005). They occur as fissural lava fields and central volcanoes installed within the main rifts and fracture zones (RP, EVZ, NVZ and TFZ; see Figure 2.2 in Volume 1). Much less frequent alkaline basalts occur in an “off-axis” position on the Snæfellsnes Peninsula (29, 30 and 31 in Figure 1.1) and on the Vestmannaeyjar Islands (Heimaey, Surtsey; 13 in Figure 1.1) located in the extension of the EVZ. Finally, basalts intermediate between the two previous types, labeled transitional basalts, form large central volcanoes in the southern part of the EVZ (Eyjafjallajökull, Katla, Tindfjöll, Hekla, Torfajökull; 14–18 in Figure 1.1), as well as, in the southeast of the island, the Öroefi volcanic chain (comprising Öræfajökull, the highest point in Iceland, Esjufjöll beneath Vatnajökull and Snæfell, north of Vatnajökull; 26 to 28 in Figure 1.1).

Figure 1.2. Icelandic magma series. Simplified diagram (modified from Jakobsson et al. 2008; Martin and Sigmarsson 2010)

These three types of basalts are distinguished by different alkali contents for the same level of SiO2 (Figure 1.2). Tholeiitic basalts are oversaturated or saturated in

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silica and alkaline basalts are undersaturated in silica; transitional basalts display intermediate geochemical and mineralogical characteristics between the two previous types (Sigmarsson and Steinthorsson 2007; Sigmarsson et al. 2008). It can be seen in Figure 1.2 that each basalt type is associated with intermediate and evolved lavas, forming magma series. These range from basalts very rich in magnesium (Mg) called magnesian basalts, picritic basalts or picrites to rhyolites. Thus, the tholeiitic series include intermediate lavas called ferrobasalts, ferroandesites and icelandites, and lead to silica-rich lavas (dacites, rhyodacites and tholeiitic rhyolites), whereas the alkaline series lead to alkaline or even hyperalkaline rhyolites (Martin and Sigmarsson 2007, 2010). However, in general, we do not observe a decrease in the proportions of the different types of lava associated with an increase in their silica content, as in most magma series resulting from fractional crystallization processes. Indeed, Iceland yields about 10% rhyolites (Walker 1965; Sæmundsson 1979) and this proportion reaches or even exceeds 20% in many central volcanoes with bimodal magma associations (basaltic and rhyolitic; Jonasson 2007; Jakobsson et al. 2008). Examples include Krafla (Nicholson et al. 1991; Jonasson 1994), Torfajökull (Gunnarsson et al. 1998), Hekla (Sigmarsson et al. 1992), Katla (Lacasse et al. 2006; Oladottir et al. 2018) and Askja (Kuritani et al. 2011), that is the peripheral volcanoes of the Icelandic hotspot. In the latter case, basaltic magmas began their evolution by assimilation coupled with fractional crystallization within a basicrustal or infracrustal magma reservoir located at a depth of about 18–20 km (0.6 GPa). Then the transition from intermediate liquids (ferrobasaltic) to tholeiitic rhyolitic magmas took place in a much shallower reservoir (0.1 GPa, about 3 km deep) (Kuritani et al. 2011). The variations of major elemental oxides as a function of MgO in young Icelandic lavas are shown in Figure 1.3. Lavas with MgO >12%, called picritic basalts or picrites, are rare; most of them result from mechanical accumulation processes of olivine phenocrysts in normal basaltic magmas (Kokfelt et al. 2006). The very strong increase in total FeO (up to 18%) and TiO2 (up to 5%) contents in the intermediate lavas of the tholeiitic series (ferrobasalts and ferroandesites) is due to the late fractionation of titanomagnetite, which itself reflects the low oxygen fugacities characteristic of the evolution of these series. Finally, the significant scatter of Na2O and K2O in the evolved lavas (MgO < 4%) is related to the differences between the alkaline and hyperalkaline rhyolites of the Snæfellsnes Peninsula and the tholeiitic rhyolites of the other volcanic centers.

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Figure 1.3. Evolution of major elements (% oxides) as a function of MgO in young Icelandic magma series (source: GEOROC database1)

The petrogenetic relations between Icelandic rhyolites and the associated basalts have been the subject of much controversy: for instance, at the beginning of the era of plate tectonics, some Icelandic geologists thought that the abundance of rhyolites showed that the substratum of Iceland was continental in nature. In fact, it can be seen in Figure 1.4 that the isotopic compositions in neodymium (Nd) and strontium (Sr) of the vast majority of Icelandic rhyolites overlap with those of the associated basalts. The alkaline and hyperalkaline rhyolites of the Snæfellsnes Peninsula, as well as some samples from Katla (Lacasse et al. 2006), are characterized by Nd isotope ratios lower than those of tholeiitic rhyolites from other central volcanoes (Figure 1.4) (Martin and Sigmarsson 2007, 2010), but identical to those of the associated alkaline basaltic magmas; their study shows that they derive from the fractional crystallization of the latter (Kokfelt et al. 2009). The tholeiitic dacites, rhyodacites and rhyolites of the central rift and fracture zone volcanoes occasionally exhibit slightly more radiogenic Sr (Figure 1.4) and lead (Pb) isotope ratios than those of the spatially and temporally associated tholeiitic basalts.

1 www.georoc.mpch-mainz.gwdg.de.

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Figure 1.4. Sr-Nd isotopic compositions of recent Icelandic basalts and associated rhyolites (simplified from Kokfelt et al. 2006; Sigmarsson et al. 2008; Martin and Sigmarsson 2010)

They are currently considered to derive from the partial melting, at the base of the Icelandic crust, of partially altered tholeiitic metabasalts (old oceanic crust). This process generated magmas of icelandite or dacite type that would have subsequently evolved by fractional crystallization (coupled with the assimilation of the wallrock basalts of the magma reservoirs) towards residual rhyolitic liquids (Martin and Sigmarsson 2007, 2010). This melting was triggered by the high geothermal gradients of the rifts, thus reproducing, all things considered, the phenomena of the genesis of acid magmas by anatexis of metabasalts that prevailed during the Archean (Martin et al. 2008; Willbold et al. 2009). Lower geothermal gradients in “off-axis” zones such as the Snæfellsnes Peninsula would inhibit this partial melting process, with fractional crystallization of alkaline basaltic magmas becoming the predominant mechanism of rhyolite genesis (e.g. for the active Ljosufjöll volcano and the Baula batholite, Figure 1.5) (Martin and Sigmarsson 2010).

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Figure 1.5. The Baula rhyolitic batholith (3 My old) at the entrance to the Snæfellsnes Peninsula (© Brigitte Van Vliet-Lanoë)

1.2.2. Geochemical diversity of young Icelandic basalts and their sources Icelandic alkaline basalts are easily distinguished from tholeiites at the major element level by their lower CaO/Al2O3 ratios at equivalent MgO content (Figure 1.6(A)), with transitional basalts (not shown in Figure 1.6) occupying intermediate fields between the two previous types. This difference reflects the fact that tholeiitic magmas evolve mainly by plagioclase fractionation (Plg in Figure 1.6(A)) whereas in alkaline magmas, clinopyroxene separation (Cpx) is dominant. Picrites result mainly from the accumulation of olivine (Ol in Figure 1.6(A)) in primitive basaltic magmas. In terms of incompatible trace elements (Figure 1.6(B)), alkaline basalts are easily distinguished from tholeiites by their greater enrichment in all of these elements, which is all the more marked as the elements considered are more incompatible (from right to left of the normalized diagram to the primitive mantle). These enrichments are similar to those of the Ocean Island Basalts (OIB, like those of St. Helena), while those of tholeiites are close to those of the rocks from oceanic ridges (MORB-N and oceanic gabbros). Fitton et al. (1997) proposed a pertinent representation of these differences using a Nb/Y versus Zr/Y diagram (Figure 1.6(C)), in which all Icelandic basalts define a linear correlation, with increasing levels of enrichment in incompatible Nb and Zr elements from picrites to tholeiites and alkaline basalts.

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Figure 1.6. Geochemical comparison of tholeiitic basalts (tholeiites), alkaline basalts and magnesian basalts (picrites) of Iceland. The fields of transitional basalts are not shown

The last two types plot globally within the OIB field, even tholeiites, the latter being more enriched than the MORB-N of oceanic ridges, which have no strict equivalents in Iceland.

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In detail, however, the slopes of the multi-elemental spectra of Icelandic basalts show variations that reflect not only the degree of enrichment of their source(s), but also the partial melting rate of the source(s), the latter being inversely proportional to the pressure exerted over the melting zone. Thus, in the case of Katla (Figure 1.7), a sharp increase in the partial melting rate occurred just after rapid deglaciation, and the early postglacial basalts emplaced at this stage were less rich in highly incompatible elements such as light rare earths (left of diagram) than the basalts of the glacial and late postglacial stages (Sims et al. 2013) (see also section 2.2.1).

Figure 1.7. Evolution of the average rare-earth composition of recent magnesian basalts from Katla (from Sims et al. 2013)

The isotopic compositions of Icelandic basalts are also heterogeneous, as shown by the relatively wide scatter of their strontium and neodymium isotopic ratios (Figure 1.4). It can be seen in Figure 1.8 that the ranges of their Nd, Hf (hafnium) and Pb isotopic ratios, although largely overlapping with that of the North Atlantic MORB, tend towards lower values in Nd and Hf. Their scatter reflects the intervention of at least three mantle components: a depleted MORB mantle (DMM) similar to the source of Atlantic MORB, and two (or more) enriched components of EM-1, EM-2 and HIMU types, related to the origin of OIB (Zindler and Hart 1986). The most 208Pb enriched compositions of Icelandic basalts also approach that of the “common C component” (Hanan and Graham

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1996), which is considered ubiquitous in the upper North Atlantic mantle (Blichert-Toft et al. 2005). The isotopic diversity of Icelandic basalts (Figure 1.8) has been the subject of much work (Fitton et al. 1997; Hanan and Schilling 1997; Hanan et al. 2000; Breddam 2002; Chauvel and Hémond 2002; Thirlwall et al. 2004; Kokfelt et al. 2006; Kitagawa et al. 2008; Peate et al. 2009, 2010). Most of these authors attribute it to the small-scale heterogeneity of the Icelandic plume, which is considered to contain enriched components in varying proportions in time and space, notably originating from slivers of subducted oceanic crust entrained within the deep mantle. Thus (Figure 1.9), Kitagawa et al. (2008) describe it using a depleted component D very close to DMM and two “local” enriched components E-1 and E-2 of the “OIB source” type; Thirlwall et al. (2004) consider four Icelandic mantle components, two of them depleted (ID1 and ID2) and two others which are enriched, IE1 and IE2 (Figure 1.9).

Figure 1.8. Compared isotopic compositions (Pb, Hf, Nd) of Icelandic and North Atlantic oceanic basalts (Agranier et al. 2005; Blichert-Toft et al. 2005)

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Figure 1.9. Lead isotopic compositions of Icelandic basalts

COMMENT ON FIGURE 1.9. – Data sources identical to Figure 1.8. Intrinsic components of the Icelandic plume: D, E-1 and E-2 from Kitawaga et al. (2008); ID-1, ID-2, IE-1 and IE-2 from Thirlwall et al. (2004). NHRL: Northern Hemisphere Reference Line. The “bulge” in 208Pb/204Pb observed for 206Pb/204Pb = 18.5 corresponds to the “anomalous” samples from Öræfajökull. The genesis of Öræfajökull transitional lavas in particular has been the subject of very divergent interpretations. They present unusual isotopic compositions, enriched in 87Sr, 207Pb and 208Pb (Prestvik et al. 2001), indicative of the contribution of an “original” enriched component to their genesis. This component is considered by most authors as mantle originating from the plume and enriched by the incorporation of subducted oceanic crust containing about 0.5% terrigenous (Prestvik et al. 2001; Peate et al. 2010) or pelagic (Kokfelt et al. 2006) sediments. It has also been proposed that these sediments originate from recycled oceanic crust present within the upper mantle underlying the volcano (Manning and Thirlwall 2013); this assumption would be consistent with the supposed occurrence of subducted oceanic plate slivers incorporated within the North Atlantic mantle (Schiffer et al. 2015).

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However, a radically different interpretation of the isotopic particularities of Öræfajökull has been proposed by Torsvik et al. (2015). These authors explain them by the assimilation of 2 to 6% of ancient continental crust coming from an isolated fragment of Jan Mayen’s microcontinent by the primitive magmas of this volcano. It is currently difficult to decide between these hypotheses, which all provide a fairly satisfactory account of the available geochemical data. 1.2.3. Some geochemical constraints concerning the origin and geodynamic evolution of Iceland To what extent can the numerous geochemical studies mentioned above contribute to a better understanding of the evolution of Iceland? The answer to this question is not obvious, as the nature and role of the mantle plume underlying the island is still under debate. Let us recall first of all (see section 3.2.5 of Volume 1) that since Wolfe et al. (1997) demonstrated the occurrence of a plume of warm mantle with a diameter of about 200 km, extending down to 400 km deep and generally consistent with Morgan’s (1971) model, all subsequent tomographic studies confirmed the existence of this plume. It has been identified within the lower mantle (Bijwaard and Spakman 1999; Montelli et al. 2006) and for some authors down to the core-mantle boundary (Helmberger et al. 1998; He et al. 2015). It has been classified among the rare plumes whose very deep origin is considered to be demonstrated (Courtillot et al. 2003; Montelli et al. 2006). It should be noted, however, that the continuity of the plume at depths greater than 600 km is debatable (Foulger et al. 2000, 2001; Hosseini et al. 2018), as the proposed deep tomographic models are highly variable (Figure 3.19 in Volume 1). On the other hand, the need for a deep plume to explain the Icelandic hotspot and associated magmatism has been questioned (Foulger and Anderson 2005). The studies dealing with the genesis of the North Atlantic Igneous Province (section 3.2.5 of Volume 1) have shown the importance of the relationships between volcanism and extensive regional tectonics in the context of rifting (Meyer et al. 2007; Hansen et al. 2009), although the hypothesis of the predominant role of an Icelandic proto-plume retains many supporters (Saunders et al. 2007). One argument in favor of the deep origin of the Icelandic plume is based on the high 3 He/4He ratios of the island’s lavas (Hanan and Graham 1996; Farley and Neroda 1998). These studies suggest the presence in their source(s) of very deep, undegassed mantle (DePaolo and Manga 2003; Class and Goldstein 2005; Graham et al. 2016; Willhite et al. 2019), although different interpretations of these ratios have been proposed (Parman et al. 2005) (section 3.2.5.3 of Volume 1).

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Similarly, the presence in the Icelandic lava mantle sources of the enriched components EM-1, EM-2 and HIMU, whose origin is generally attributed to the contribution of slivers of oceanic crust subducted at very great depths (Zindler and Hart 1986), is consistent with a very deep origin. Indeed, the systematic presence of these end-members in lavas from hotspots linked to deep plumes (Courtillot et al. 2003; DePaolo and Manga 2003), has allowed many authors (references in (Peate et al. 2010)) to model the genesis of island lava by mixing enriched materials from a very heterogeneous plume and North Atlantic-type upper mantle (depleted DMM and enriched C components). It is necessary to take into account, at least locally, other components of crustal or mantle origin. These include, in many cases, more or less altered basaltic materials of the Icelandic crust, which are involved either in the processes of assimilation coupled with the fractional crystallization of basaltic magmas, or as a source of evolved lavas (section 1.2.1). In addition, isotopic peculiarities of Öræfajökull lavas could be due to either their contamination by an underlying ancient microcontinental fragment (Torsvik et al. 2015), or to the presence within the upper North Atlantic mantle of components inherited from subducted oceanic plates (Manning and Thirlwall 2013). In summary, geochemistry allows the detection of the presence in lava sources of crustal or mantle components whose nature remains identifiable, but it is much more difficult to position them from a geodynamic point of view. 1.3. Central volcanoes and active fissural systems 1.3.1. Central volcanoes The construction of a large volcanic structure generates, on the surface, tensions in the upper crust of a magnitude comparable to that of regional tectonic constraints and to that of the overpressure occurring within the magma chamber. The maximum stress is reached at the top of the chamber, at the apex of the volcanic structure. This implies a preferential fragility of the chamber walls along this axis and thus the focusing of the volcanic activity at the level of a central crater (Pinel and Jaupart 2003), often leading to the formation of a caldera. At the level of the shallow magma reservoir underlying the shield volcano or stratovolcano, the magmatic overpressure at the beginning of the eruption increases as the structure grows, but decreases following its eventual destruction. Simple volcanoes, of interglacial type, are mostly strictly effusive shield volcanoes such as Ok or Skjaldbreiður, characterized by the potential existence of lava lakes and post-emissive collapse cauldrons (Figure 1.10). These volcanoes represent early manifestations of deglaciation: the great majority of them were formed

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a little before 8,000 years BP, from large subcrustal reservoirs fed by unloading during deglaciation. They may be reactivated during successive interglacials, such as in the case of Þeistareykir. They are at the origin of very large channeled lava fields with numerous tunnels (decimetric to decametric in diameter), similar to those of Hawaiian volcanism or extraterrestrial volcanism (Mars and Venus). The high fluidity of the lavas suggests a main period of emplacement just after the last deglaciation, unlike fissural volcanoes (> 8,000 years BP).

Figure 1.10. Examples of Icelandic shield volcanoes

COMMENT ON FIGURE 1.10. – (A) Skjaldbreiður, an interglacial shield volcano, located north of Lake Þingvellir (10.3 ky?) (© Brigitte Van Vliet-Lanoë). (B, C) Shield volcano Þeistareykir (Jökulsá á Fjöllum), active during the deglaciation of the Bølling and probably Termination II (transition MIS 6a and MIS 5e): summit crater and collapsed lava tunnels. Same scale; line length: 100 m (© LMIs). The central volcanoes located at the apex of the hotspot generally display a relatively continuous, mainly basaltic activity, even during the Ice Age (Grímsvötn, Bárðarbunga, Kverkfjöll). Peripheral volcanoes such as Torfa, Katla, Askja, Öræfa, Hekla, Krafla and Kerlingarfjöll have a mixed basaltic and rhyodacitic activity, related to the Plio-Quaternary neo-rift (about 3 My, see Chapter 2 of Volume 1) and glacial dynamics. Generally, these large central volcanoes are linked to a fissure swarm. In the

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Eastern Volcanic Zone, the very large Torfa caldera and the Katla caldera have an estimated age of 700 ky. The base of the Oræfajökull dates back to 1.5 My. Many rhyolitic volcanoes were emplaced or had a strong activity around 400 ky, related to the overload of the largest quaternary cap. This is also the case of Hekla, whose oldest known activity dates back to 413 ky (Guillou et al. 2010). They are sometimes associated with a dome-shaped morphology, similar to that of volcanoes on Venus (Figure 1.11). Other types of central volcanoes may have a dominant activity during interglacials, but also as subglacial edifices. This is for instance the case of Askja, a volcano with a multiphase caldera. The volcanic activity is then dual, with relatively “dry” explosive phases of strombolian type alternating with phreatomagmatic phases (section 1.4). A given eruption, depending on the accumulation of water in the caldera, involves either abrupt or gradual changes between plinian and/or phreatomagmatic explosions and effusive activities. The classical sequence is an initial plinian rhyolitic plume followed by the emplacement of basaltic tephras. It is often linked to the primary emptying of an evolved crustal reservoir, later refilled from an infracrustal reservoir (Hartley and Þórðarson 2013). In periods of deglaciation, eruptive phases may be more frequent. The March 1875 (rhyodacitic) eruption of Askja was the consequence of a major volcano-tectonic episode (1874–1876) in the Northern Volcanic Zone. Rift-related tectonic activity commonly interferes with the activity of volcanoes such as Torfa (Figure 1.11). Finally, some central volcanoes such as Hekla are essentially of fissural type.

Figure 1.11. Rhyolitic volcano (on the right; diameter 250 m) with dome-like morphology. The yellow star indicates its location within the Torfajökull caldera in the EVZ, Reykjadalir (LMIs images)

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1.3.2. Fissural volcanism and subaerial lava flows 1.3.2.1. Modes of emplacement In Iceland, rift-related activity is characterized by effusive volcanism emitted along active fissures en échelon, linked to extension (see Chapter 2 of Volume 1). These fractures may be either in connection with one or more deep (infracrustal) reservoirs linked to the central volcanoes, as is the case with Hekla, Krafla or Bárðarbunga (10% of eruptions) (Larsen and Guðmundsson 2014), or may be located within the rift downstream from these central volcanoes.

Figure 1.12. Large Icelandic fissural eruptions. (A) Lakagígar fissure, 1783–1784 (© Hervé Guillou). (B) Veiðivötn fissure, 1744 (© Ingibjorg Kaldal). (C) Holhuraun eruption, 2014–2015 (© Reykjavik Helicopter)

The compositions of the emplaced lavas are mainly controlled by the deep reservoirs (section 1.2). Aerial eruptions always begin with the outpouring, from lava fountains, of scoriae that weld together when they fall back around the eruption site, and of fluid lava flows emitted from the entire fault length or from the

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en échelon fault network. Then they concentrate by ejecting gases from the high points of the fissure system, generating welded scoriaceous cones referred to as spatter cones, or even temporary lava lakes such as those of Holhurhaun, while downstream lava outpouring continues. For fissure swarms connected to a central volcano, the eruption begins with intense seismic activity at the level of the volcano, or even with an explosive eruption, depending from the volume of lava injected into the shallow magma reservoir. This was the case during the eruption of the “Mývatn fires” of 1724–1729. This eruption began with the formation, from May 17 to 18, 1724, of an explosive crater on the southern flank of Krafla, the Viti crater. It then continued with the outpouring of a 34 km long lava flow reaching Mývatn Lake. For the recent eruptive phase of Krafla (1975–1984, eight successive eruptions), it was the seismic crisis of December 20, 1975 that initially opened fissures following the inflation of the magma chamber located at a shallow depth (3 km) beneath the caldera and caused the introduction of a 19 km long lava flow sequence. An additional W-E extension of 9 cm was measured during this event. In the case of the last fissural eruption of Bárðarbunga in mid-August 2014, it was also a two-week seismic crisis associated with the emptying of the lava reservoir located beneath the caldera (injection of the Dyngjuháls dyke, active up to the eastern base of Askja) that preceded the fissure eruption of Holhuraun (Figure 1.12(C)), which continued until May 2015. Aerial photographs show similar eruptive scenarii during the Holocene. The Veiðivötn fissural eruption in 1477, of the same type, was curiously associated with a pyroclastic eruption almost synchronous with that of Torfajökull (4 months later), connected to another magma reservoir in the southward extension of the same rift. In this case, fissural volcanism had, in part, a phreatomagmatic evolution due to the subsident character of the rift, located at the western edge of Vatnajökull and characterized by a high regional water table. The Grímsvötn-Katla rift presents the same type of activity as the Veiðivötn rift; this activity is mostly located at the edge of the ice cap. This was the case with the eruptions of the Eldja in 932–941 AD and the Laki (Lakagígar) in 1783–1784. The Eldja eruption, even more voluminous than that of the Laki, was probably initiated by Katla activity in the south (tephra recorded in Greenland ice); it emitted 20 km3 of lava just after the Viking settlement (see Chapter 3). It also emitted 220 Mt of SO2, and was associated with the emission of jökulhlaups (subglacial discharges) at the level of the Mýrdal glacier’s sandur. The eruption of the Laki emplaced 14.7 km3 of basaltic lavas during seven months of activity. It was fed by the Grímsvötn volcano, and culminated with the outpouring of the giant Skaftáreldar flow (Figure 1.13).

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Figure 1.13. Map of the fissural eruptions of the Eldgjá and Laki (Southeast Neorift) and their lava flows (modified from Sigurðardóttir 2014)

During late-glacial periods, this fissural magmatism essentially formed hyaloclastite ridges aligned over fault swarms. Sometimes, these ridges emerged and then emplaced subaerial volcanics, often of phreatomagmatic type. These networks of faults reactivated during unloading by deglaciation. Especially, a particularly powerful fissural eruption occurred around 11–10.8 ky, from Askja towards the coast at Melrakkaslétta, hitting the North-East Volcanic Zone sideways (Figure 1.14). At that time, the glacier front was located south of the Hrossaborg tuff cone with a much lower topographic slope than the present one, partly flooded by numerous lakes. This eruption was associated with a major rhyolitic tephra, the Askja S, most probably related to a reactivation of the caldera of the same name, and several phases of lava outpouring (Sigurgeirsson 2016). The fissural eruptions of the Askja-Bárðarbunga-Veiðivötn group (Askja S and flows from Þjórsárdalur, Figure 1.14) marked major emissions in the EVZ at the end of a climate-induced deglaciation, but prior to the major isostatic rebound of

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Vatnajökull; this could be a consequence of the unloading at the top of the Icelandic hotspot. The feeding of large dykes (> 20 m in width) at the origin of these fissural eruptions was often triggered, during deglaciation, from the subcrustal reservoir zone, due to the mantle unloading during deglaciation (see section 2.1). These injections generally occurred during rifting episodes forced by the crustal extension unlocked by the glacioisostatic discharge, common during deglaciation.

Figure 1.14. (A) Askja S fissural eruption, 11–10.8 ky, (modified from Sigurgeirsson 2016). Black dots: cones and volcanoes; purple: lava flows. (B) Spatter cone (Rauðholar). (C, D, E) Pseudocraters and lava lake injected under light gray tillite of Hljóðaklettar (white arrows, Vestudalur, Jökulsá á Fjöllum) (photos © Brigitte Van Vliet-Lanoë)

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1.3.2.2. The diversity of subaerial lava flows Lava flow emplacement in Iceland is mainly controlled by fault swarms, located either on rift faults or on radial faults, the latter allowing the magma column in a central volcano like the Bárðarbunga to be drained. The volume of the flows is directly related to the piezometric level of the magma above the magma chamber(s). The typology of flows is mainly morphological (Figures 1.15 and 1.16). The viscosity of the flows, usually related to their speed, determines the dominant surface morphology. Fluid lava with a ropy surface (pahoehoe flows) and viscous lava with a brecciated surface (aa flows) have a very different appearance in the field, but their geochemical composition can be identical or very similar. A single flow sometimes displays a pahoehoe surface upstream, aa downstream, and becomes completely brecciated when it stops its progression (block flow or block lava).

Figure 1.15. Examples of postglacial basaltic flows. (A) Southwest of Lake Hagongulὀn. Note the digitation of the flows and the feeding tubes in relief. (B) Transition between a smooth ropy lava and a valley drain displaying an aa facies (Reykjanes). (C) Very fluid lava spread in a maze of tubes (< 5 m in diameter): postglacial Snaefells cone (LMI images)

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Fluid lava flows are often emplaced by successive pulses during the same eruption. Magma is channeled within lava tubes fed by the continuation of the emission under the solidified crust. Their morphology is often digited, when they are about 10 meters thick, sometimes leaving depressions between the lava fingers. If they are flowing on a slope, their surface morphology is uneven with fractured domes called tumuli (Figure 1.16) and “pressure ridges” elongated more or less perpendicular to their propagation axis. Tumuli are sometimes emplaced in an underwater environment and are then identical to certain forms observed at the level of oceanic ridges (Deschamps et al. 2016).

Figure 1.16. Morphology of fluid basaltic flows. (A) Ropy lava surface (“Mývatn fires”). (B) smooth polygonal lava fractured and reinjected (Jökulsá to Fjöllum, slope towards Askja). (C) Morphology in plurimetric tumuli (Reykjanes peninsula), with (D) a massive lava tore at its core (© Brigitte Van Vliet-Lanoë)

During the deglaciation period, the volume of fluid magma available at the edge of the ice cap was significant: the associated flows were often very extensive, not very thick (a few meters) and invaded the valleys or run down the cliffs. They filled up topographic depressions, forming thin to decameter-thick lava lakes, with a surface that was often smooth and polygonal or roped (pahoehoe) when emplaced over a slight slope. The result was a morphology quite rarely observed on land, such as in Dimmuborgir, near Lake Mývatn (Figure 1.17), with towers showing traces of ancient

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lava lake levels and depressions paved with large fragments of roped lava. These morphologies are nearly identical to those observed under deep water at the level of oceanic ridges. Brecciated surface aa type flows are much more common during the Holocene and interglacials than pahoehoe flows, and their composition may be basaltic, ferrobasaltic, ferroandesitic, icelanditic, rhyodacitic or rhyolitic (Figure 1.18).

Figure 1.17. Morphology and post-eruptive evolution of the Dimmuborgir lava lake (Lake Mývatn, ca 2,000 years BP, resulting from a fissural eruption, the Lúdentsborgir) with (A) roof collapses (modified from the Natural park scheme), (B, C, D) lava tunnels and traces of successive slipped lava crusts, (E) steam vents generally constituting the residual pillars (© Brigitte Van Vliet-Lanoë)

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Figure 1.18. Morphology of viscous rhyolitic (obsidian) lava of aa type. Landmanalaugar (image (A) © LMIs; photos (B, C, D) © Brigitte Van Vliet-Lanoë)

Figure 1.19. Vesicles and lava flows. (A) Pipe vesicles at the base of a fluid lava flow from the top of a hyaloclastite tuya (Eiriksjökull). (B) Vesicles at the base of a tumulus formed by an underwater lava flow (Reykjanes) (© Brigitte Van Vliet-Lanoë)

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Figure 1.20. Examples of lava flows

COMMENT ON FIGURE 1.20. – (A) Ropy lava flows (80 m high; south coast of the Reykjanes peninsula) (© Brigitte Van Vliet-Lanoë). (B) Superposition of two types of basaltic lava flows: at the base, composite flows formed by a number of successive cooling units; at the top, vertical jointing of individual lava flows separated by scoriaceous breccia (© René Maury). Cliffs of Svörtuloft, Snæfellsnes peninsula (height 15 m). The thickness of these lava flows can reach more than 100 meters, and their surface may be spiked with decametric blocks, especially in the case of highly viscous rhyodacitic or rhyolitic acid lavas (Landmanalaugar and Hekla obsidians). The internal structure of basaltic lava flows is clearly visible at the level of coastal

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cliffs and rift escarpments. Under their smooth, ropy or brecciated surface, the individual lava flows are generally massive with rare gas vesicles, especially at their base (Figure 1.19), and present a clearly visible vertical columnar jointing (Figure 1.20) in the Icelandic Neogene lava piles. On the other hand, the composite lava flows, which are emplaced as successive lava streams, occur as juxtaposed cooling units of metric to decametric size, whose coarse radial jointing is reminiscent of that of pillow lavas. 1.3.3. Hydromagmatism Hydromagmatic eruptions occur when rising magma meets relatively shallow water: ocean, lake or subglacial lake, phreatic water table (the latter case generating phreatomagmatic eruptions). The magma, ejected at more than 1,000°C, degasses suddenly, forming cypressoid plumes of very black ash. The lava explodes due to the pressure reached by the gases and water vapor or simply by thermal shock, breaking it up into fragments of submillimetric to decimetric sizes by coalescence of the vesicles and tempering of the glass. These debris accumulate as hydrated and yellowish altered glass fragments known as palagonitic, gray peperites in a marine context (reducing, alkaline environment) or hyaloclastites, which become ochre by oxidation in a partially subaerial environment. Hyaloclastites can only form under relatively low pressure, when the height of the water column above the magma is less than about 700–1,000 m. If the gas pressure is low, pillow lavas are emitted at different depths. Under atmospheric pressure, they can form in the immediate vicinity of the coast, at the apex of lava tubes, as it is currently the case in Hawaii. Under a glacier, the pressure will vary depending on whether the supravolcanic aquifer is confined or not. Pillow lavas are generally basaltic or intermediate, rarely rhyolitic. They are formed during eruptions with a relatively low rate of effusion in a cold aqueous environment, and are fed either by a dyke or a lava tube. The slow extrusion of lava at a temperature of 1,000°C to 1,200°C allows the instantaneous formation of rods or elongated tores with a flexible glassy skin that thickens around each individual lava lobe, preventing their coalescence. Pillows detach under the effect of gravity and accumulate in cones or plurimetric heaps, sometimes intermixed with hyaloclastites, in shallow water environments (Figure 1.21). Under low pressure, the pillow lavas are rich in vesicles arranged in concentric zones. Under deep subglacial conditions, they generally contain none.

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Figure 1.21. Examples of pillow lavas. (A) Pillow lavas included within hyaloclastites, tuya of Litlasaltvík, south of Húsavík. (B) Pillow lavas without a matrix, base of the Helludindar ridge (Kleifarvatn, Reykjanes) (© Brigitte Van Vliet-Lanoë)

1.3.3.1. Submarine or Surtseyan volcanism The first indication of an ongoing submarine volcanic eruption is a persistent plume of steam on the ocean surface and sometimes the presence of gas bubbles and floating pumice rafts. Then appear the first explosions in cypressoid sprays (Figure 1.22) that burst the ocean surface and end up building a cone of tuffs that emerges progressively, like that of Surtsey Island in 1962–1963 south of the Vestmann Islands, an offshore extension of the Eastern Neo-volcanic Zone (Katla, Oræfajökull).

Figure 1.22. Cypressoid hydromagmatic plumes, loaded with black hydrated ashes, Surtsey, 1962 (© J.S. Aabech, Norway)

These islands are fragile and do not resist well to the action of swell and storms. Once the island emerges, classical eruptive processes, often of fissural type,

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take over with an explosive Strombolian-type volcanism (scoriae and lapilli cones), eventually followed by the emplacement of lava flows that consolidate the island, as was the case in Heimaey in 1973.

Figure 1.23. Eruptions of the Grímsvötn volcano. (A) Melting cauldron. (B) Melting lake. (C) Hydromagmatic eruptive plume. (D) Microscopic thin-section view of the hydromagmatic tephra (Grímsvötn 127 ky, scale 50 µm): note the hydrated orange rim of the fragments (palagonite) (photos A, B, C © ISOR; photo D © Brigitte Van Vliet-Lanoë)

1.3.3.2. Sub-lacustrine or sub-glacial hydromagmatism In Iceland, the emplacement of hyaloclastites is commonly observed in subglacial position, and is most often associated with a meltwater lake. The internal pressure of the gases in the magma must be in equilibrium with the pressure exerted by the overlying ice cap, generally of the order of 600 to 700 m of ice. Hyaloclastites are often interstratified with turbiditic lacustrine facies and present progradations related to the accumulation of fine debris around the emission point. When the ice overload disappears following an eruption, the pressure decreases and pillow lavas are often observed interstratified with the hyaloclastites. These rules apply to the formation of hyaloclastite ridges (palagonite ridges) and that of sub-glacial volcanoes. At present, the height of ice above subglacial meltwater lakes

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is about 400 m (Bjornsson 2017). During an eruption, the ice cap collapses in the lake and forms a cauldron (Figure 1.23). 1.3.3.3. Phreatomagmatism Phreatomagmatism is very frequent during deglaciation. Indeed, during glacio-isostatic rebound, the slope of the valleys is temporarily softened due to the rapid rebound of the coastal zone, limiting runoff and facilitating marine transgression on the still depressed substratum. Meltwater is therefore abundant (ice caps, periglacial lakes, melting permafrost). Phreatomagmatism can take different forms. Phreatomagmatic tuff rings, such as the Hverfjall in the Mývatn basin (Figure 1.24) or the Hrossaborg in the Jökulsá valley, Fjöllum, are rooted on a lavaemitting fault that came into contact with a high phreatic water table (proximity of the glacial front or location in a graben).

Figure 1.24. Examples of phreatomagmatic deposits. (A) Hverfjall tuff cone (Mývatn, 2500 BP (for Before Present, before 1950)). (B) Accretionary lapilli also known as “volcanic oolites”, formed by the aggregation of fine ashes to water droplets. (C) Dropstone impact pattern projected into water-saturated tephra (line scale: 10 cm) (© Brigitte Van Vliet-Lanoë)

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Figure 1.25. Examples of emplacement of phreatomagmatic formations

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COMMENT ON FIGURE 1.25. – (A) Rootless lapilli cones, Mývatn (ca 2,500 years BP) formed over the Hverfjall lava flow. (B) Þjorsardalur lava flow from a fissural eruption, with phreatic degassing pipes. (C) Pile of phreatomagmatic bombs with chilled margins. (D) Diagram of the formation of Þjórsárdalur rootless cones (adapted from (Reynolds et al. 2005)) (photos A, B, C © Brigitte Van Vliet-Lanoë). These well-layered tuffs are made up of hydrated glass, most often oxidized, sometimes aggregated in the form of small spheres (accretionary lapilli) and presenting many bomb impact figures or dropstones (Figure 1.24). The most frequent forms are tuff or lapilli cones without roots with summital pseudocraters as in the case of Lake Mývatn (Figure 1.25). A field of such rootless cones occupies the Þjórsárdalur, where a subaerial fissural flow has covered a former coastal plain. These volcanic features are most often emplaced during the effusion of lava over waterlogged surfaces (sediment or aquifer in fractured rock), or even in lake depressions as in Mývatn. 1.4. Volcanic hazards in Iceland 1.4.1. Hazards related to lava flows Fluid lava flows are relatively harmless in Iceland. Their most negative impact is usually a significant SO2 degassing (see section 1.5.2). Nevertheless, several recent cases have interfered with human activities. These are the Laki flows in 1783, the Krafla flows in 1724–1729 and 1975–1984, and the eruption of Eldfell (Heimaey Island) in 1973. Laki fissural lava flows followed the river beds, submerging farm after farm for the first 12 days. Krafla is a very active fissural volcano, due to its location at the intersection of the Northeast Fault System (NVZ) with the Tjörnes Transform Zone (TFZ), which is connected to the Kolbeinsey Oceanic Ridge (KR). The 1724–1729 “Mývatn fires” lavas flowed from the Viti Crater and stopped on the north shore of the present Mývatn lake after submerging two farms and the houses of the parish of Reykjahlíð. During the activity phase of 1975–1984, the lava flows stopped near the pre-existing geothermal power plant, downstream in the valley. On the island of Heimaey (Vestmann Islands, south of Iceland), the fissural eruption of the Eldfell volcano in January 1973 produced a lava flow that threatened the opening of the island’s natural harbor. The lava produced was diverted and cooled by spraying seawater, which allowed the formation of a peninsula, thus increasing the protection of the harbor area (Figure 1.26(B)). Another technically

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sensitive area is the Hengill power plant, which supplies the Reykjavik area with hot water and electricity, as well as the Straumur aluminum plant. The last eruption of the Hengill in this area dates back to 2,000 years BP.

Figure 1.26. The 1973 Eldfell eruption

COMMENT ON FIGURE 1.26. – (A) The two main stages of the Eldfell eruption illustrated by Icelandic stamps: lava fountains forming spatter cones along the new fissure, followed by the growth of the Eldfell breached cone which emitted a flow towards the port (the hill on the right is Helgafell). See also the introduction to this volume. (B) The growth of Heimaey Island during the eruption, according to a USGS document. 1.4.2. Hazards related to explosions and gas emissions Iceland explosive volcanism is mainly characterized by hydromagmatic eruptions, largely emitting black ash or basaltic tephra, often followed by Strombolian explosions. The volumes of ash emitted can be enormous (> 450 km3) as in the case of the serial eruptions of Saksunarvatn tephra (10.3–9.9 ky), a volcanic crisis linked to the deglaciation of the Vatnajökull area and involving at least the

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Grímsvötn and Bárðarbunga volcanoes (Guðmundsdóttir et al. 2016). This tephra, which was scattered over an area of 1,500,000 km2 and found as far as continental Europe, constituted a remarkable chronological line for the stratigraphy of the North Atlantic and that of Greenland ice cores. Another pyroclastic volcanism with a cyclicity of 500 to 1,000 years is documented from the central volcanoes peripheral to the hotspot, independently of deglaciations. It is an evolved volcanism of rhyolitic to rhyodacitic nature (section 1.2), characteristic of an inter- to fini-glacial, but especially interglacial context. The eruptions are mainly of Plinian type, with emission of a stratospheric plume of gas and gray fine ashes that will drift with the atmospheric circulation, as well as pyroclastic flows involving base surges with more local but catastrophic effects. Often, this type of eruption evolves in a second stage towards a basaltic eruption of phreatomagmatic type. The most violent eruption was that of the Katla, located under the Mýrdalsjökull ice cap. It marked the beginning of Holocene deglaciation: the Vedde tephra or Vedde Ash (Figure 1.27) is in fact dated between 12 and 11.8 ky BP (Lane et al. 2012).

Figure 1.27. Extension of Vedde tephra in the North Atlantic (stars) and adjacent continents (modified and completed from Tomlinson et al. 2011)

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The Hekla volcano is responsible for one of the largest rhyolitic eruptions of the Icelandic Holocene, that of the tephra Hekla 3 (or H3) dated around 1000 BCE, which projected about 7.3 km3 of volcanic rock into the atmosphere and caused an 18-year-long climatic cooling. The Hekla eruption of 1693 (Volcanic Explosivity Index or VEI = 4) was one of the most destructive eruptions in history and lasted for seven months, emplacing 0.18 km3 of tephra, covering the northwestern part of the island, causing lahars and flash floods (local glaciers), and destroying many farms and wooded moors in Þjórsárdalur. The ashes reached Norway, and resulted in a significant destruction of the terrestrial and aquatic fauna there via emitted chlorine and fluorine.

Figure 1.28. Examples of rhyolitic deposits. (A) Subaerial pumice deposit (Hekla, H3 3,000 BP). (B, C) Ignimbrite (thermally welded tuff) from Þórsmörk, Torfajökull 57 ky. (D) Base surge deposit from a pyroclastic flow emitted by Askja in 1875 (© Brigitte Van Vliet-Lanoë)

The rhyolitic pyroclastic eruption of 1362 AD in Öræfajökull (Volcanic Explosivity Index or VEI = 5), of which 2.3 km3 of tephra are preserved, probably released about 10 km3; its modeling shows that a similar future eruption would cover the eastern half of Iceland with tephra, causing major economic and communication problems (Barsotti et al. 2018). The most recent rhyodacitic

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pyroclastic eruption in Iceland was that of Askja in 1875 (Figure 1.28). The Icelandic population lives away from active volcanoes and pyroclastic flows. The only volcano relatively close to inhabited areas is Hekla, whose eruptions are often forecasted less than two hours before they start. The ash and gas eruptions, on the other hand, with a much more distal impact, are at the origin of very significant and often insidious damages.

Figure 1.29. Examples of tephras. (A) Spreading of fine ash from the Eyjafjallajökull (© Veðurstofa Íslands). (B) Glass fragment rolled by a jökulhlaup in a lake sequence (Grímsvötn). (C) Pumice fragment from the Vedde tephra (Katla). (D) Basaltic glassy splinters (Grímsvötn) (photos B, C, D, scanning electron microscope © Brigitte Van Vliet-Lanoë)

A first impact is a cooling of the climate: volcanic ashes propelled into the stratosphere and picked up by jet-type currents dim sunlight for several months or years, regardless of their geochemical nature. The gaseous emissions of SO2 evolve very quickly into sulfuric aerosols, which reflect sunlight. Basaltic eruptions emit large quantities of gases, especially SO2, chlorine and fluorine, which are toxic for fauna and humans. In addition, tephras are also the cause of respiratory accidents

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due to their splintered structure (Figure 1.29) and their fixation on moist mucous membranes and in the lungs. The fissural eruption of Laki in 1783 devastated Iceland and buried the summer pastures under lava flows and especially ashes. It produced more than 120 million tons of SO2. Humans and cattle were starved and poisoned by the “hunger fog” during the winter of 1783–1784 and the cold years that followed: this volcanic fog was indirectly responsible for the death of 10,000 Icelanders, a fifth of the island’s population (Jackson 1982). This lethal spiral was exacerbated by the intrinsic poverty of the island and its economic isolation. The prevailing winds blew south and east and exported these gases directly to Europe (Þórðarson and Self 2003), especially during the summer of 1783. This dry, acidic fog reached London, Paris, Stockholm, Rome and beyond, causing respiratory problems. On a more global scale, the loading of the air with sulfuric aerosols considerably reduced insolation. Temperatures cooled throughout the northern hemisphere, and the winter of 1783–1784 was one of the coldest in the Little Ice Age (Grattan 2005). The climate was then disrupted for years, reducing the flood flow of the Nile and causing famine in Egypt. It is estimated that the direct and indirect effects of Laki’s eruption killed more than one million people worldwide in the 18th century. The fissural eruption of Eyjafjallajökull in 2010 (Figure 1.30) was not powerful by Icelandic standards, but it had a considerable economic impact on the rest of Europe (Karlsdóttir et al. 2012).

Figure 1.30. Eruption of Eyjafjallajökull in April 2010. (A) Eruptive plume of the eruption (image Veðurstofa Íslands). (B) Modeled ash concentration over Europe (modified from Stuefer et al. 2013)

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As in the case of the eruption of Laki in 1783, the prevailing winds blew south and east, directly into European airspace. Tephra particles suspended in the upper atmosphere were abrasive and dangerous to aircraft engines because of their fusibility at relatively low temperatures. As a result, civil aviation officials had to make the decision to block almost all commercial flights over the Atlantic and Western Europe, but Akureyri airport in the north of the island remained operational. The current management of eruptive risks in Iceland has taken into account the impact of these two eruptions whose frequency is of the order of 10 years for common eruptions and of the century for major eruptions, all volcanoes combined. 1.4.3. Jökulhlaups and associated hazards Ice breakup or glacial flooding is one of the powerful and spectacular processes that occur at the end of glaciation at the expense of meltwater damming lakes.

Figure 1.31. Evolution of the lake level in the Grímsvötn caldera between 1930 and 2000. The ice in the crater rose until a jökulhlaup drained the lake, leading to the formation of a cauldron (Figure 1.23). In 2000, the lake level also caused the ice dam to rise and float, allowing a major flood to be discharged (according to Bjornsson 2000)

Ice dam breaks are the most common generative phenomenon. Simple climatic meltwaters have a fairly high viscosity due to their low temperature. In most cases, the level of the lake supplied with meltwater will rise (Figure 1.31) until the glacier body acting as a dam is lifted along the flow path to allow the water to pass through; it may even explode and break up the glacier surface into icebergs. In this case, the frequency of water discharges increases linearly, culminating in a time interval of

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several hours to 1–2 days without conduit enlargement by melts. The eventual drainage of an aquifer fed by surface water infiltration and flooding the glacier crevasse network may sustain a high flow for several weeks or even months. In Iceland, the great majority of jökulhlaups, these sudden subglacial and/or juxtaglacial breakup events (Figure 1.32), emerge at the surface or downstream, in relation to subglacial volcanic activity (formation of a volcano or a hyaloclastic ridge, birth of a geothermal field). Nevertheless, supraglacial or subglacial lakes can also be fed by the infiltration of surface waters.

Figure 1.32. Some effects of jökulhlaups

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COMMENT ON FIGURE 1.32. – (A) Traces of the resurgence of a jökulhlaup in the Grìmsvötn caldera, in 2000: rupture of the glacier surface. (B) Spreading of icebergs and their dispersal on the downstream Skeiðarárjökull sandur during the same event. (C) Esker of jökulhlaup drainage, Kringslá river, Brúarjökull. (D) Field view of iceberg melting traces. (E) Glacial flood moraine (surge), with basal shear level. (F) Melt subsidence (kettle hole) linked to the melting of an iceberg in a preboreal jökulhlaup (11.4 cal ky; Ljósavatn school) (photos A, B © Tom Van Loon; C: LMIs image; D, E, F © Brigitte Van Vliet-Lanoë). Drainage of the epi-volcanic lake occurs by gravity at low pressure, less than the static ice overload. Volcanic meltwater is relatively warm and triggers the digging of tunnels or eskers, diverting the water along the edge of the ice cap. The conduits, often networks of crevasses in the ice or hydraulic fractures at the top of the substratum, develop slowly over days or weeks and then drain abruptly. The heat input being maintained by the volcanic activity, the hydrostatic pressure on the walls is cyclically rebalanced, generating a pulsating behavior at the water outlet. All intermediate processes can occur. In general, a simple jökulhlaup episode generates a subglacial eruption by discharge, as in the case of the Kverkfjöll volcano in northeast Iceland (Höskuldsson et al. 2006); or the reverse occurs by spreading of ashes on the surface of the glacier and subsequent acceleration of surface melting linked to the decrease in albedo (Wittmann et al. 2017). Normally, jökulhlaups do not lead to glacial floods, but the hydraulic overpressure generated by eruptions of glacier-covered volcanoes can cause a sudden detachment of the basal surface of the glacier and retroactively trigger a glacial rapid advance or surge, or even the form hydraulic explosion of the ice front as in 2000. The corresponding deposits have a shallow terminal moraine (< 10 m) characterized by an organization in stacked folds with imbricated blocks, less spectacular than those of ice river floods (see Chapter 3). 1.4.3.1. Morphological and sedimentological impacts Many Icelandic canyons were formed during these recurrent catastrophic floods. They are indeed likely to incise canyons in the bedrock within a few days and to introduce deposits of turbulent flows at high pressure. The largest jökulhlaups (107 m3) recorded in Iceland were created by the emptying of calderas, lakes or subglacial aquifers. The emergent waters are often heavily loaded with sediments, boulders and often ice blocks. The presence of melt holes or kettle holes in the sedimentary facies is often a sign of the presence of icebergs or hydraulic dismantled glaciers.

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Initially, the charged fluid behaves as a non-Newtonian (non-compressible) liquid, like a mudflow, and the flow becomes cohesive or even supercritical over large areas. Subglacial lake jökulhlaups can transport about 107 tons of sediment per event, but during violent volcanic eruptions, the sediment load discharged has reached a total of 108 tons. Powerful tractive phenomena are responsible for “tail of comet” accumulations, often parallel, of plurimetric blocks. Megadunes of plurimetric size can also form in the flow constriction zone. Jökulhlaups have a significant erosive potential (Figure 1.33): they excavate large canyons over time and transport and deposit huge quantities of sediments and icebergs (Figure 1.34) towards the coastal plains or sandur. The flow rocky surfaces can be dismantled and the volcanic cones truncated. The jökulhlaups are responsible for a rapid blunting of the blocks and an oriented but less regular polish than that produced by wind abrasion.

Figure 1.33. Examples of jökulhlaup-related erosion

COMMENT ON FIGURE 1.33. – (A) Jökulsá á Brú canyon downstream of the Hálslón Dam (80 m high). The other photos represent the Jökulsá á Fjöllum valley: (B) Asbyrgi

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dead end (the road gives the scale); (C) Hafragil dead end and canyon (Dettifoss); (D) oriented polish by the sand loaded waters of jökulhlaups (Kreppa; height 1 m) (photos: A © Hervé Guillou; B, C, D © Brigitte Van Vliet-Lanoë).

Figure 1.34. Examples of jökulhlaups-related sedimentary deposits

COMMENT ON FIGURE 1.34. – (A) Jökulhlaup deposits, Ljósavatn (school). (B) Stacking of plurimetric blocks, Jökulsá á Fjöllum valley. (C) Spreading of megadunes (300 m wide) of Bárðarbunga basaltic tephra (10.1–9.9 ky BP) in Hólssandur, Jökulsá á

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Fjöllum (the road on the left gives the scale). (D) Gravel megadunes (width 50 m: the road gives the scale; Jökulsá á Fjöllum). (E) 20 m high megadune, Vesturdalur, Jökulsá á Fjöllum. (F) Erosion in pots, pillars and blades, Vikingavatn (Öxarfjördur) in relation to a turbulent flow with cavitation (photos A, B, F © Brigitte Van VlietLanoë; C, D, E © LMIs). 1.4.3.2. Retroaction on volcanism These drainages can lead by unloading to a local decompression of the shallow magma reservoirs, in the same way as deglaciations and episodes of global warming, and thus retroactively induce volcanic eruptions that will, in turn, recharge the subglacial waters and spread volcanic ash. This was the case during the complex Grímsvötn-Bárðarbunga eruption from 10.3 to 9.9 ky BP (Saksunarvatn tephra), during the final Preboreal deglaciation. Jökulhlaups resulting from subglacial volcanic eruptions were quite brutal and dangerous, and evolved into lahars and debris-laden floods. 1.4.3.3. Jökulhlaups and climate The significant jökulhlaups in Iceland are related to the presence of ice caps and deglaciation events. The majority of jökulhlaups studied occurred along the Jökulsá á Fjöllum (Kverkfjöll and Bárðarbunga volcanoes; summary in (Baynes et al. 2015)) in the north, and the Skeiðarársandur (Grímsvötn) in the south, all originating from Vatnajökull and the outlets of the Mýrdalsjökull (Katla volcano) like those from Markarfljὀt. In fact, these outlets have been functional since the major glaciations (3 My; Chapter 3). The Þjórsá and Rangá valleys in the south, and those of the Jökulsá á Fjöllum and Skjálfandi in the north preserve the testimony of the passage of recurrent jökulhlaups during the Eemian interglacial (Van Vliet-Lanoë et al. 2018). Other indications of jökulhlaups exist in relation to the HofsjökullKerlingarfjöll glacial complex and the Langjökull ice cap (Figure 1.35). Their occurrence was mainly recorded during the Younger Dryas, and up to 8.2 ky BP, in association with melting episodes connected with Bond events (ice calves in the North Atlantic), a Holocene variant of the Dansgaard–Oeschger (DO) events of the Ice Age without an ice shelf. Floods disappeared during Thermal Optimum, then reappeared with glaciers from the Subboreal (< 5 ky BP) and during the Middle Age thermal Optimum (Van Vliet-Lanoë et al. 2020b). A series of floods marked the warming after the Little Ice Age, in relation to the glacial floods of 1880.

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Figure 1.35. Map of the extension of postglacial jökulhlaups (maximum flood influence; completed from Van Vliet-Lanoë et al. 2020b))

1.4.3.4. Jökulhlaups and humans As early as the year 1000 AD, jökulhlaups threatened farms and pastures; the agricultural decline in the Þjórsá and Ytri Rangá valleys, attributed to Hekla eruptions, such as those of 1104 (H1), 1158 and 1693 AD, was also due in part to major erosion of soil and grazing land by floods. Most of the abandoned farms in Þjórsárdalur had lost most of their pasture (Figure 1.36). Old developments in floodplains show that jökulhlaups have damaged grassland areas, disrupted roads in the sandur, changed property boundaries, and even caused flood waves in coastal waters. In Iceland, most of the energy production is hydroelectric, with dams and power plants installed on the drainage channels of large ice caps. Currently, most of the meltwater collection facilities are dams built with a technical zone of structural weakness, allowing limited building failure (earth dams, made waterproof with a concrete coating), and quick and much cheaper repairs. Nevertheless, these buildings, planned for jökulhlaups with centennial return time, are undersized by a factor of 10 compared to the estimated volumes of the largest Holocene jökulhlaups.

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The same is true for the single-track bridges, which are mainly built with steel girders and beams to be easily and economically replaced.

Figure 1.36. Traces of soil erosion (highlighted by vegetation) induced by jökulhlaups in the Þjὀrsá valley (top, glacial gray water)

COMMENT ON FIGURE 1.36. – The red dots correspond to abandoned medieval farms. The flow direction is towards the left. Note the transverse walls at the eroded areas aimed at reducing the speed of the current (18th century; spacing about 150 m) (photo LMIs). 1.4.4. Icelandic dust: a consequence of volcanism Iceland is the most active volcanic center in the North Atlantic. It is, along with Jan Mayen Island, the source of most of the tephra and dust emissions found in North Atlantic sea cores and in the peat bogs of Europe, and even in the ice cores of Greenland. Iceland is also a land particularly subject to winds, especially from the northeast, in relation to the Icelandic depression, located mostly in the southwest of the island. As a result, the island is the locus of intense wind activity.

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The fixation of eolian sediments by the capillary humidity of soils, wetlands, snows or even mosses or cryptogamic crusts in arid contexts is, by far, the most important mechanism. This fixation concerns both wind-tracted and suspended particles and allows their capture. Synchronous rise of the capillary fringe, as the deposit settles, maintains the fixation. On the other hand, the presence of a substrate that is often basaltic, and therefore fractured, limits the “fixing” role of the water table’s capillary fringe. 1.4.4.1. Tephras Iceland is one of the countries where tephrostratigraphy was born, particularly in the northern and southern central areas of the island (Þórarinsson 1944, 1970). No less than 10 Icelandic volcanic centers are tephra emitters. Tephrochronology appears as a tool allowing both precise dating and inter-regional synchronization of past volcanic events. Ejecta (tephras or pyroclastites) are made up of fragments of magma expelled into the air during the eruption of a volcano and solidified at a given moment of the eruption, or generally during their aerial trajectory (see section 1.3). The term tephra is generally used for fine ash. Tephra is most often glassy, but can include feldspars and other rock fragments torn from the volcanic chimney. Proximal deposits can be quite coarse (lapilli), but those dispersed over long distances by eruptive plumes are generally very fine (2–10 µm). Tephras can be directly sedimented in a sequence, most often continental: they are then primary deposits, as is the case in Icelandic loess or Greenlandic ice cores, with 100% of the initial geochemical characteristics of the tephra preserved or only modified by a possible meteoric alteration. Modest alteration or mixing may result from runoff or wind action penecontemporary with the deposit (Figure 1.37). They may be reworked during large fluvial incisions and then resedimented in a more recent estuarine formation (Van Vliet-Lanoë et al. 2018). They may also drift by floating on the surface of lakes or the ocean or even on ice rafts; they will then settle in secondary sediments such as sea cores, but are not free of elutriation or mixing. Tephras are also very sensitive to frost and easily reduced to silty particles during cold periods (seasonal or of glacial rank). Tephras are materials that are difficult to date, especially by the 40Ar/40K or Ar/40Ar methods (see section 2.3), because of their origin: a mixture of magma and vent sweeping products from the volcanic chimney, when they do not derive from a complete explosive destruction of the structure. The global analysis of a tephra may eventually enable defining its volcanic source, but certainly not the genuine chemical composition of the eruption.

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Figure 1.37. (A) Deflation by north wind overflowing the sea, in the extension of the Markarfljót valley (south of the island, Earth Observatory image, Nasa). Note in A the algal production in turquoise blue fertilized by dust (2 swirls). (B) Deflation in the Jöklusá valley at Fjollum (© Brigitte Van Vliet-Lanoë)

Figure 1.38. Main Late glacial and Early Holocene tephras in relation to jökulhlaups (arrows) and climate (NGRIP curve)

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COMMENT ON FIGURE 1.38. – PBE: Preboreal cooling; AL: Alleröd; BCE: intraBølling cooling. Note their unusual concentration from 10.4 ky, the date corresponding to the beginning of the Saksunarvtan (Norway) tephra events (modified from Van Vliet-Lanoë et al. 2020b). Nevertheless, a geochronological approach is possible, especially on potassium feldspar, for local eruptive tephras whose geochemistry attests to their juvenile character. The dating of tephras is generally established indirectly (Figure 1.38), either by marine or organic materials that frame the deposit, or by an accumulation age model for ice cores. The characterization of tephras is done by geochemical analysis at the grain scale using electron microprobe for major elements and laser ablation ICPMS for trace elements. Grímsvötn tephras very often have the same geochemical signature, including trace elements (Oladottir et al. 2011). The vesicle morphology of pyroclasts can also be indicative of the source. Hekla pumices are characterized by vesicles of heterogenous sizes. Those of Katla (Vedde Ash) are much more regular. Geochemistry is not the only criterion for dating a tephra: during the Late Glacial, Katla produced at least three rhyolitic tephras with the same geochemical signature between 15 ky and 9 ky. The Skógar tephra is a supraglacial mixture of these different Katla tephras, which led to stratigraphic confusions. A rhyolitic tephra from the Askja volcano, Askja S or Askja 10 (11.4 cal ky), with splintered morphology, is subcontemporary to the Katla tephras (Van Vliet-Lanoë et al. 2020b). Generally, large tephra emissions occur during deglaciation episodes, when the volcanic structure is still covered by an ice cover (< 300 m). Tephrochronology enables calibrating for recent periods (post 18th century) other dating methods such as lichenometry or to date glacial advances. Tephras, especially if they are basaltic and easily weathered, bring silica, iron and phosphates to the oceanic masses, thus increasing the fertility and fish richness of North Atlantic waters (Zimbelman and Gregg 2000; Prospero et al. 2012) in addition to the direct alluvial inputs from the island (Figure 1.37(A)). Significant algal developments appear around Iceland during periods of volcanic activity, eolian activity or important contributions from rivers (summer melting or jökulhlaup). 1.4.4.2. Loess An important characteristic of Iceland is the frequency of loess deposition on either side of the central zone of the island. Their composition is 100% volcaniclastic, either fresh (tephra) or residual, the latter resulting from hydrothermal alterations or fluvio-glacial spreading. In Iceland, loess is interstratified with numerous tephras

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(Figure 1.39). They are not uniformly distributed in the landscapes, but are mostly accumulated on the cold or humid northern and western slopes, whereas they have been eroded or stretched by solifluction on the south-facing slopes. Holocene loess can reach 7–8 m thick in the southeast between the Mýrdal and Vatna glaciers and 2 m on the slopes of Eyjafjörður (around Akureyri) or Vopnafjörður. While some of them date from the Allerød and the beginning of the Preboral, the majority was emplaced around 8.5 ky BP (Jackson et al. 2005), after the actual end of deglaciation.

Figure 1.39. Interstratified rhyolitic and basaltic tephras (A) in the Vopnafjörður peat bogs (northeast) and (B) in the loess of the Ytri Rangá (south) (© Brigitte Van Vliet-Lanoë)

This important part of wind sedimentation is indicative of seasonal fluvio-glacial activity, in a context where soil erosion is effective (ice needles or pipekrakes, runoff), but where and when aridity takes over, especially at the end of summer. It is the signature of cooling associated with a deterioration of the plant cover after a “hotter” and wetter episode. The spreading of tephra, by smothering the vegetation, also

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promotes this deflation. The impact of jökulhlaups is especially sensitive by the local destruction of loessic or sandy soils and their resedimentation downstream. Icelandic loess are very sensitive to runoff and suffosion (formation of underground runoff tunnels), allowing the shaping of morphology into very spectacular gullies in the south of the island (Figure 1.40(C)) and wind erosion patterns, in combination with frost, rain and runoff. This activity is partly maintained by sheep browsing, forming rofaborðs (Icelandic term characterizing an erosion morphology with the overhanging herbaceous stratum; Figure 1.40(D)). They are linked to episodes of climatic degradation at the beginning and end of the Holocene, in association with a lowering of the water table and the formation of loessic hummocks. The latter are emblematic of Iceland, and are referred to as thufurs. They can appear in three to four years by frost heave of the soil in winter, with even greater effectiveness when the loess cover is thin, particularly on basalt. Many of these loess contain leaf debris (polar willow, dryad) and therefore are not indicative of a very arid environment.

Figure 1.40. Examples of Icelandic loess deposits and their erosion

COMMENT ON FIGURE 1.40. – (A) Deposit of thick recent loess on exposed slope north of the Landmanalaugar Valley, interstratified with tephra. (B) Thin section cut in

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recent loess with a basaltic component (Jökulsá á Fjöllum; line length: 50 µm). (C) Erosion gullies in loess deposits south of Laki. (D) Residual “rofaborð” morphology, Jökulsá á Fjöllum Valley, linked to combined erosion by wind and sheep activities (photos © Brigitte Van Vliet-Lanoë). From an economic point of view, loess is one of the most useful substrates: the soils are very fertile, basaltic glass being highly susceptible to weathering, and its water retention capacity being high, particularly due to the formation of amorphous clays of the allophane or imogolite type. They are therefore the optimal substrate of grassy or even cultivated soils. Nevertheless, overgrazing promotes wind erosion and contributes to the destruction of vegetation cover in addition to the inland clearing of land since colonization. 1.4.4.3. Continental dunes Dune sedimentation is carried out according to several classical parameters of wind dynamics. A dynamic of traction and saltation of the sandy fraction leads to the constitution of flat ridges (2D transverse dunes) at moderate wind speed (large laminar flow), barkhan dunes (3D dunes) at higher speed and parabolic dunes evolving into longitudinal dunes at speeds below 120 km/h in the presence of turbulent flows. In the absence of surface roughness (vegetation, blocks or micro-relief), the strong winds (> 150 km/h) allow the installation of flat laminates (Figure 1.41(C)), or sand sheets, characteristic of the proximity of glacial fronts. Dry, flat surfaces offer very little wind resistance and do not promote major deposition. A dynamic of fixation by the roughness of the substrate (swirls, loss of competency) leads to the trapping of the fractions pulled or salted by herbaceous and bushy vegetation (Figure 1.41(A)) or by obstacles, such as in sub-desert or coastal areas. The majority of Icelandic and South-Greenlandic dune fields are of this type. The deposits can be interbedded with snow (Figure 1.41(D)) and collapse following thaw. Sandurs and especially the areas affected by jökulhlaups constitute the major sand resource of Iceland (Figure 1.41(B)). Winnowing ridges appear on the surface of recent volcanic ash deposits, such as near Askja or on the western flanks of the Hekla.

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Figure 1.41. Examples of Icelandic dunes. (A) Vegetated dunes (Grímsstaðir). (B) Vegetated dunes reworking the surface of jökulhlaup hydraulic dunes (northwest Hrossaborg, LMIs image). (C) Laminated wind sand (Grímsstaðir). (D) Wind-blown sand deposit (Jökulsá á Fjöllum) (A, C, D © Brigitte Van Vliet-Lanoë)

1.5. References Baynes, E.R.C., Attal, M.L., Niedermann, S., Kirstein, L.A., Dugmore, A.J., Naylor, M. (2015). Erosion during extreme flood events dominates Holocene canyon evolution in northeast Iceland. PNAS, 112(8), 2355–2360. Breddam, K. (2002). Kistufell: Primitive melt from the Iceland mantle plume. J. Petrol., 43, 345–373. Campbell, I.H. (2007). Testing the plume theory. Chem. Geol., 241(3–4), 153–176. Chauvel, C., Maury, R.C., Blais, S., Lewin, E., Guillou, H., Guille, G., Rossi, P., Gutscher, M.A. (2012). Marquesas isotopic stripes constrain the size of plume heterogeneities. Geochem. Geophys. Geosyst., 13(1) [Online]. Available at: http://doi.org/10.1029/ 2012GC004123. Class, C. and Goldstein, S.L. (2005). Evolution of helium isotopes in the Earth’s mantle. Nature, 436, 1107–1112. Courtillot, V., Davaille, A., Besse, J., Stock, J. (2003). Three distinct types of hotspots in the Earth’s mantle. Earth Planet Sci. Lett., 205, 295–308.

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DePaolo, D.J. and Manga, M. (2003). Deep origin of hotspots – The mantle plume model. Science, 300, 920–921. Farley, K.A. and Neroda, E. (1998). Noble gases in the Earth’s mantle. Ann. Rev. Earth Planet. Sci., 26, 189–218. Fitton, J.G., Saunders, A.D., Norry, M.J., Hardason, B.S., Taylor, R.N. (1997). Thermal and chemical structure of the Iceland plume. Earth Planet. Sci. Lett., 153, 197–208. Graham, D., Michael, P., Shea, T. (2016). Extreme incompatibility of helium during mantle melting: Evidence from undegassed mid-ocean ridge basalts. Earth Planet. Sci. Lett., 454, 192–202. Grattan, J. (2005). Pollution and paradigms: Lessons from Icelandic volcanism for continental flood basalt studies. Lithos, 79, 343–353. Gunnarsson, B., Marsh, B.D., Taylor Jr., H.P. (1998). Generation of Icelandic rhyolites: Silicic lavas from the Torfajökull central volcano. J. Volcan. Geotherm. Res., 83, 1–45. Hanan, B.B. and Graham, D.W. (1996). Lead and helium isotope evidence from oceanic basalts for a common deep source of mantle plumes. Science, 272, 991–995. Hanan, B.B. and Schilling, J.G. (1997). The dynamic evolution of the Iceland mantle plume: The lead isotope perspective. Earth Planet. Sci. Lett., 151, 43–60. Hanan, B.B., Blichert-Toft, J., Kingsley, R., Schilling, J.G. (2000). Depleted Iceland mantle plume geochemical signature: Artifact of multicomponent mixing? Geochem. Geophys. Geosystems, 1 [Online]. Available at: http://doi.org/10.1029/1999GC000009. Hartley, M. and Maclennan, J. (2018). Magmatic densities control erupted volumes in Icelandic volcanic systems. Front. Earth Sci., 6(29) [Online]. Available at: http://doi. org/10.3389/feart.2018.00029. Hartley, M.E. and Þórðarson Þ. (2013). The 1874–1876 volcano-tectonic episode at Askja, North Iceland: Lateral flow revisited. Geochem. Geophys. Geosys., 14, 2286–2309. He, Y., Wen, L., Capdeville, Y., Zhao, L. (2015). Seismic evidence for an Iceland thermochemical plume in the Earth’s lowermost mantle. Earth Planet. Sci. Lett., 417, 19– 27. Helmberger, D.V., Wen, L., Ding, X. (1998). Seismic evidence that the source of the Iceland hotspot lies at the core-mantle boundary. Nature, 396, 251–255. Höskuldsson, A., Sparks, R.S.J., Caroll, M.R. (2006). Constraints on the dynamics of subglacial basalt eruptions from geological and geochemical observations at Kverkfjöll, NE-Iceland. Bull. Volcan., 68, 689–701. Jackson, E.L. (1982). The Laki eruption of 1783: Impacts on population and settlement in Iceland. Geography, 67, 42–50. Jackson, M.G., Oskarsson, N., Trønnes, R.G., McManus, J.F., Oppo, D.W., Grönvold, K., Hart, S.R., Sachs, J.P. (2005). Holocene lœss deposition in Iceland: Evidence for millennial scale atmosphere-ocean coupling in the North Atlantic. Geology, 33(6), 509–512.

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Jakobsson, S.P., Jonasson, K., Sigurdsson, I.A. (2008). The three igneous rock series of Iceland. Jökull, 58, 117–138. Karlsdóttir, S., Petersen, G.N., Björnsson, H., Pétursson, H., Þorsteinsson, H., Arason, Þ. (2012). The Eyjafjallajökull eruption, 2010 – The role of IMO [Online]. Available at: http://en.vedur.is/earthquakes and volcanism/articles/nr/2072. Kokfelt, T.F., Hoernle, K., Hauff, F., Fiebig, J., Werner, R., Garbe-Schonberg, D. (2006). Combined trace element and Pb, Nd, Sr, O isotope evidence for recycled oceanic crust (upper and lower) in the Iceland mantle plume. J. Petrol., 47, 1705–1749. Kokfelt, T.F., Hoernle, K., Lundstrom, C., Hauff, F., Van den Bogaard, C. (2009). Timescales for magmatic differentiation at the Snæfellsjokull central volcano, western Iceland: Constraints from U-Th-Pa-Ra disequilibria in post-glacial lavas. Geochim. Cosmochim. Acta, 73, 1120–1144. Kuritani, T., Yokohama, T., Kitawaga, H., Kobayashi, K., Nakamura, E. (2011). Geochemical evolution of historical lavas from Askja Volcano, Iceland: Implications for mechanisms and timescales of magmatic differentiation. Geochim. Cosmochim. Acta, 75, 570–587. Lacasse, C., Sigurdsson, H., Carey, S.N., Johannesson, H., Thomas, L.E., Rogers, N.W. (2007). Bimodal volcanism at the Katla subglacial caldera, Iceland: Insights into geochemistry and petrogenesis of rhyolitic magmas. Bull. Volc., 69, 373–399. Lane, C.S., Blockley, S.P.E., Mangerud, J., Smith, V.C., Lohne, Ø.S., Tomlinson, E.L., Matthews, I.P., Lotter, A.F. (2012). Was the 12.1 Ka Icelandic Vedde Ash one of a kind? Quat. Sci. Rev., 33, 87–99. Larsen, G. and Guðmundsson, M.T. (2014). Volcanic system: Bárðarbunga system [Online]. Available at: https://m.en.vedur.is/media/jar/Bardarbunga_kafli20140825.pdf. Manning, C.J. and Thirlwall, M.F. (2013). Isotopic evidence for interaction between Öræfajökull mantle and the Eastern Rift Zone, Iceland. Contrib. Mineral. Petrol., 167 [Online]. Available at: http://doi.org/10.1007/s00410-013-0959-1. Martin, E. and Sigmarsson, O. (2007). Crustal thermal state and origin of silicic magmas in Iceland: The case of Torfajökull Ljosufjöll and Snæfellsjökull volcanoes. Contrib. Mineral. Petrol., 153, 593–605. Martin, E. and Sigmarsson, O. (2010). Thirteen million years of silicic magma production in Iceland: Links between petrogenesis and tectonic setting. Lithos, 116, 129–144. Martin, E., Martin, H., Sigmarsson, O. (2008). Could Iceland be a modern analogue for the Earth’s early continental crust? Terra Nova, 20, 463–468. McGarvie, D.W., Burgess, R., Tindle, A.G., Tuffen, H., Stevenson, J.A. (2006). Pleistocene rhyolitic volcanism at the Torfajökull central volcano, Iceland: Eruption ages, glaciovolcanism, and geochemical evolution. Jökull, 56, 57–75. Meyer, R., Van Wijk, J., Gernigon, L. (2007). The North Atlantic Igneous Province: A review of models for its formation. Geol. Soc. Am. Spec., 430, 525–552.

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Montelli, R., Nolet, G., Dahlen, A., Masters, G.A. (2006). Catalogue of deep mantle plumes: New results from finite-frequency tomography. Geochem. Geophys. Geosyst., 7 [Online]. Available at: http://doi.org/10.129/2006GC0001248. Morgan, W.J. (1971). Convective plumes in the lower mantle. Nature, 230, 42–43. Nicholson, H., Condomines, M., Fitton, J.G., Fallick, A.E., Grönwold, K., Rogers, G. (1991). Geochemical and isotopic evidence for crustal assimilation beneath Krafla, Iceland. J. Petrol., 32, 1005–1020. Óladóttir, B., Larsen, G., Sigmarsson, O. (2011). Holocene volcanic activity at Grímsvötn, Bárdarbunga and Kverkfjöll subglacial centres beneath Vatnajökull, Iceland. Bull. Volcan., 73, 1187–1208. Óladóttir, B., Sigmarsson, O., Larsen, G. (2018). Tephra productivity and eruption flux of the subglacial Katla volcano, Iceland. Bull. Volcan., 80(58) [Online]. Available at: http://doi.org/10.1007/s00445-018-1236-y. Parman, S.W., Kurz, M.D., Hart, S.R., Grove, T.L. (2005). Helium solubility in olivine and implications for high 3He/4He in ocean island basalts. Nature, 437(7062), 1140–1143. Peate, D.W., Baker, J.A., Jakobsson, S.P., Waight, T.E., Kent, A.J.R., Grassineau, N.V., Skovgaard, A.C. (2009). Historic magmatism on the Reykjanes Peninsula, Iceland: A snap-shot of melt generation at a ridge segment. Contrib. Mineral. Petrol., 157, 359–382. Peate, D.W., Breddam, K., Baker, J.A., Kurtz, M., Barker, A.K., Prestvik, T., Grassineau, N., Skovgaard, A.C. (2010). Compositional characteristics and spatial distribution of enriched Icelandic mantle components. J. Petrol., 51, 1447–1475. Pinel, V. and Jaupart, C. (2003). Magma chamber behavior beneath a volcanic edifice. J. Geophys. Res., 108(B2), 2072 [Online]. Available at: http://doi.org/10.1029/2002JB001751. Prospero, J.M., Bullard, J.E., Hodgkins, R. (2012). High-latitude dust over the North Atlantic: Inputs from Icelandic proglacial dust storms. Science, 335, 1078–1082. Reynolds, R.J., Brown, T., Thordarson, T., Llewellin, E.W., Fielding, K. (2015). Rootless cone eruption processes informed by dissected tephra deposits and conduits. Bull. Volcan., 77(72) [Online]. Available at: http://doi.org/10.1007/s00445-015-0958-3. Sæmundsson, K. (1979). Outline of the geology of Iceland. Jökull, 29, 7–28. Saunders, A.D., Jones, S.M., Morgan, L.A., Pierce, K.L., Widdowson, M., Xu, Y.G. (2007). Regional uplift associated with continental large igneous provinces: The roles of mantle plumes and the lithosphere. Chem. Geol., 241, 282–318. Schiffer, C., Balling, N., Jacobsen, B.O., Stephenson, R.A., Nielsen, S.B. (2014). Seismological evidence for a fossil subduction zone in the East Greenland Caledonides. Geology, 42(4), 311–314. Sigmarsson, O. and Steinthorsson, S. (2007). Origin of Icelandic basalts: A review of their petrology and geochemistry. J. Geodyn., 43, 87–100.

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Sigmarsson, O., Condomines, M., Fourcade, S. (1992). A detailed Th, Sr, and O isotope study of Hekla: Differentiation processes in an Icelandic volcano. Contrib. Mineral. Petrol., 112, 20–34. Sigmarsson, O., Maclennan, J., Carpentier, M. (2008). Geochemistry of igneous rocks in Iceland: A review. Jökull, 58, 139–160. Sims, K.W.W., Maclennan, J., Blichert-Toft, J., Mervine, E.M., Blusztajn, J., Grönvold, K. (2013). Short length scale mantle heterogeneity beneath Iceland probed by glacial modulation of melting. Earth Planet. Sci. Lett., 379, 146–157. Storey, M., Duncan, R.A., Tegner, C. (2007). Timing and duration of volcanism in the North Atlantic Igneous Province: Implications for geodynamics and links to the Iceland hotspot. Chem. Geol., 241, 264–281. Stuefer, M., Freitas, S.R., Grell, G., Webley, P., Peckham, S., McKeen, S.A., Egan, S.D. (2013). Inclusion of ash and SO2 emissions from volcanic eruptions in WRF-Chem: Development and some applications. Geosci. Model Dev., 6, 457–468. Thirlwall, M.F., Gee, M.A.M., Taylor, R.N., Murton, B.J. (2004). Mantle components in Iceland and adjacent ridges investigated using double-spike Pb isotope ratios. Geochim. Cosmochim. Acta, 68, 361–386. Tomlinson, E.L., Þórðarson, Þ., Müller, W., Thirlwall, M., Menzies, M.A. (2010). Microanalysis of tephra by LA-ICP-MS. Strategies, advantages and limitations assessed using the Thorsmörk ignimbrite (S. Iceland). Chem. Geol., 279, 73–89. Torsvik, T.H., Amundsen, H.E.F., Tronnes, R.G., Doubrovine, P.V., Gaina, C., Kuznir, N.J., Steinberger, B., Corfu, F., Ashwal, L.D., Griffin, W.L., Werner, S.C., Jamtveit, B. (2015). Continental crust beneath southeast Iceland. PNAS, 112(15), E1818–E1827. Van Vliet-Lanoë, B., Bergerat, F., Allemand, P., Innocent, C., Guillou, H., Cavailhes, T., Liorzou, C., Grandjean, P., Passot, S. (2020a). Tectonism and volcanism enhanced by deglaciation events in southern Iceland. Quat. Res., 94, 94–120 [Online]. Available at: http://doi.org/10.1017/qua.2019.68. Van Vliet-Lanoë, B., Knudsen, O., Guðmundsson, A., Guillou, H., Chazot, G., Langlade, J., Liorzou, C., Nonnotte, P. (2020b). Volcanoes and climate: The triggering of Preboreal jökulhlaups in Iceland. Int. J. Earth Sci., 109, 847–876 [Online]. Available at: http://doi.org/10.1007/s00531-020-01833-9. Walker, G.P.L. (1965). Acid volcanic rocks in Iceland. Bull. Volcan., 29, 375–402. Willbold, M., Hegner, E., Stracke, A., Rocholl, A. (2009). Continental geochemical signatures in dacites from Iceland and implications for models of early Archaean crust formation. Earth Planet. Sci. Lett., 279, 44–52. Willhite, L.N., Jackson, M.G., Blichert-Toft, J., Bindeman, I., Kurtz, M.D., Halldórsson, S.A., Harðardóttir, S., Gazel, E., Price, A.A., Byerly, B.L. (2019). Hot and heterogenous high 3He/4He components: New constraints from Proto-Iceland plume lavas from Baffin Island. Geochem. Geophys. Geosys., 20, 5939–5967.

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Wittmann, M., Groot Zwaaftink, C.D., Steffensen Schmidt, L., Guðmundsson, S., Pálsson, F., Arnalds, O., Björnsson, H., Thorsteinsson, T., Stohl, A. (2017). Impact of dust deposition on the albedo of Vatnajökull ice cap, Iceland. Cryosphere, 11, 741–754. Zimbelman, J.R. and Gregg, T.K.P. (2000). Environmental Effects on Volcanic Eruptions: From Deep Oceans to Deep Space. Kluwer Academic/Plenum Publishers, New York. Zindler, A. and Hart, S.R. (1986). Chemical geodynamics. Ann. Rev. Earth Planet. Sci., 14, 493–571. Þórarinsson, S. and Sæmundsson, K. (1979). Volcanic activity in historical times. Jökull, 29, 29–32. Þórðarson, Þ. (2003). Laki-Grímsvötn eruptions II: Appraisal based on contemporary accounts. Jökull, 53, 11–48. Þórðarson, Þ. and Höskuldsson, A. (2002). Iceland. Terra Publishing, Dunedin Academic Press, Liverpool, Edinburgh. Þórðarson, Þ. and Larsen, G. (2007). Volcanism in Iceland in historical time: Volcano types, eruption styles and eruptive history. J. Geodyn., 43, 118–152. Þórðarson, Þ. and Self, S. (2003). Atmospheric and environmental effects of the 1783–1784 Laki eruption: A review and reassessment. J. Geophys. Res., 108(D1), 4011.

2

Volcanism and Glaciations: Forcings and Chronometers Hervé GUILLOU, René MAURY and Brigitte VAN VLIET-LANOË

2.1. Subglacial volcanic landforms One of the characteristics of the Quaternary evolution of Iceland is that the volcanic structures, currently constituting the high points of the island, allowed the construction of glaciers that evolved into ice caps. Subglacial volcanism is a form of hydromagmatism specifically induced by the melting of this ice cover. A thick ice cap covering Iceland partially limits by its loading effect the partial melting of the underlying mantle, and thus the volcanic activity, except at the level of very large volcanoes (for example, the Grímsvötn caldera had tephra emissive activity throughout the last glaciation). The classification of subglacial landforms built by volcanism is subject to much debate. We will limit ourselves here to the generic terms of tuyas (Russell et al. 2014) for roughly circular shapes, the Icelandic móbergs, fed by a relatively point source, and ridges or tindar for elongated landforms built on rift or fracture zones (see Chapter 1). The surfaces that stake the tuyas can be successive surfaces for stabilizing their growth according to the variations in ice cap thickness, but can also result from

For color version of the figures in this chapter see www.iste.co.uk/vanvliet/iceland2.zip Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021. Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

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successive “deltaic” accumulations of hyaloclastites emitted from the eruptive chimney into the lake (Smellie et al. 2006). When tuya formation occurs in a cold glacier (ice temperature < – 0.5°C), lava emission is constrained by limited ice melting, resulting in a tower-like morphology topped by hyaloclastites (Figure 2.1), and possibly followed by subaerial construction (Smellie et al. 2018).

Figure 2.1. Sub-lacustrine bedded hyaloclastites of the last glacial episode, Snæfellsnes (© Brigitte Van Vliet-Lanoë)

2.1.1. Subglacial isolated volcanoes or tuyas Tuyas (or móbergs) are the subglacial equivalent of subaerial shield volcanoes (Werner et al. 1996) and are formed during one or more subglacial eruptions, from a deep reservoir, generally in relation to fault clusters. The volume introduced during a single eruption can be very large, as in the case of the Heirðubreið. The Snæfell (northeast of Vatnajökull), on the other hand, was progressively emplaced during 100 ky. From bottom to top, tuyas are composed of pillow lavas and then of hyaloclastites surrounded by volcanic breccias. This succession indicates a decrease in external pressure within the subglacial cavity. The latter is often occupied by a lake of thermal melt water, whose size increases when the volcano grows during its development towards the surface of the ice cap, sometimes forming lateral deltas. As the ice thins, phreatomagmatism takes over pillow eruption, and hyaloclastic deposits (Figure 2.2) are emplaced.

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Figure 2.2. Hyaloclastite facies. (A) Unoxidized hyaloclastite facies (Kleifarvatn, deglaciation dated back to 155 ky) with (B) turbiditic figures and (C) figures of debris flow progradation related to the functioning of the subglacial mouth. (D, E) oxidized hyaloclastite facies (© Brigitte Van Vliet-Lanoë)

Their shape is roughly circular with steep sides and a flat top (Figure 2.3). At their base, their diameters vary from a few hundred meters to a few kilometers. Their height is generally between 200 and 1,000 m.

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Thus, the height of these tuyas can be a good indicator of the thickness of the ice cap at the time of their emplacement (Bourgeois et al. 1998; Werner et al. 1999; Smellie 2000). Once the tuya emerges from the cap, very fluid lava flows build up a shield volcano.

Figure 2.3. Examples of tuya-type recent subglacial volcanoes

COMMENT ON FIGURE 2.3. – (A) Prestahnúkur (rhyolitic tuya of Langjökull). (B, D) Herðubreið: basaltic tuya with late effusive subaerial cone (emersion of the ice cap; note the slope of the subglacial plateau). (C) Basaltic tuya: aerial view of the summit of Bæjarfjell (northwest of Krafla) with subsidence due to the emptying of a lava lake, without subaerial effusive cone (photos A and B © Brigitte Van Vliet-Lanoë; photos C and D © LMIs). When the magma activity is sufficiently intense and sustainable, horizontal subaerial lava flows can cover the tops of these volcanoes. If the glacier is of the “cold” type, the tuya has very steep slopes (Figure 2.4) and a relatively narrow diameter, due to the small volume of the subglacial lake, as in the case of the Ásgarðsfjall (Kerlingarfjöll).

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Figure 2.4. Examples of steeply sloping tuyas. (A) Diagram of a tuya with subglacial lake deltaic terraces (modified from Smellie 2000). (B) Rhyolitic tuya of the Ásgarðsfjall (Kerlingarfjöll), formed under full glacial conditions with a cold base glacier and a small melting cavity (© Brigitte Van Vliet-Lanoë)

2.1.2. Hyaloclastite ridges or tindar These ridges (Figure 2.5) were formed essentially during glaciations, over very large dykes (≥20 m wide) derived from the fault clusters of the Icelandic rifts (Figure 2.6). They are the subglacial counterpart of subaerial fissural eruptions

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(Guðmundsson 1986). They are generally derived from infracrustal basaltic reservoirs (see Chapter 1). The maximum magma supply is fed bottom up during periods of partial deglaciation or disappearance of the cap. It is generally accepted that the volumes of lava emitted during subglacial fissural eruptions are lower than those that produced tuyas (Guðmundsson 1986, 1987).

Figure 2.5. Hyaloclastite ridges. (A) To the immediate north of Vatnajökull, about 150 ky old (Kreppa). (B) Fjallgarður Ridge, with late subaerial volcanism, 400 ky old (© Brigitte Van Vliet-Lanoë)

The ridge sequences begin with thick pillow lava sequences, overlain by hyaloclastites, often stratified near the ground surface, and end with phreatomagmatic tuffs (Figure 2.7). They can be up to 40 km long, 2 to 4 km wide and

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several hundred meters high. These morphologies are often much more elongated than the central volcanoes, attesting to a multi-kilometer widening of rifts during the glacial period (Bourgeois et al. 2000), but not necessarily inherited from the last glaciation (Figure 2.5).

Figure 2.6. Formation of fissural rift volcanoes during an interglacial period (example from the Askja area) and a deglaciation period (modified from Guðmundsson 1987; Hartley and Þórðarson 2013)

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Figure 2.7. Structuration of a hyaloclastite ridge (Kleifarvatn). (A) Interpretative diagram (explanations in the text; modified from Van Vliet-Lanoë et al. 2019). (B) Rather massive pillow lava edifice at the bottom of the ridge. Note on the right the local feeding dyke (Kleifarvatn) (© Brigitte Van Vliet-Lanoë)

A particular facies, the “laminated hyaloclastites” (Smellie 2008), is present at the level of ridges formed at the immediate edge of the ice cap. These are lateral injection facies under a relatively thin ice cover, with a basal sill that may be overlain by pillow lavas, the latter being covered in turn by hyaloclastites showing injection figures of the underlying magma and indications of lateral displacements (Figure 2.8). This process generally occurs repeatedly at the edge of the cap. Good

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examples can be observed at the foot of Katla and further east, along the south coast of the island.

Figure 2.8. Facies of laminated hyaloclastites at the edge of ice sheets. (A, B) Keldunúpur (Siða-Fljótshverfi, southern district of Iceland): water/hyaloclastites/ magma interaction with basaltic injections within unconsolidated hyaloclastites. (C) Stöng (Þὀrsardalur): hyaloclastite ridge bordering the Hekla and crosscut by numerous dykes (© Brigitte Van Vliet-Lanoë)

2.2. Volcanism, deglaciation and climate 2.2.1. General features: deglaciation, discharge and partial melting Numerous studies describe the links between deglaciation and increased volcanic activity. 2.2.1.1. The facts Accounts of the history of volcanic eruptions show that subaerial volcanism increased worldwide during the last deglaciation, between 12 ky and 7 ky. This

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phenomenon occurred mainly in regions undergoing deglaciation, as did seismic activity. In southern Alaska, volcanic activity increased at the beginning of the Bølling interstadial (14,500 years BP) in relation to a strong global warming (Praetorius et al. 2016). It has been estimated for Iceland that eruption rates were 30–50 times higher during the first 1,500 years following the last deglaciation (12,000 years) than those observed during other periods (MacLennan et al. 2002), and that about 64% of the total volume of volcanic products were emitted during the first 3,000 years after deglaciation (Sinton et al. 2005). A similar phenomenon occurred after the Subboreal “cold snap” (5,500 years BP or Neoglacial) and after the Little Ice Age. Icelandic glaciers (Chapter 3) are clearly different from the Greenland and Scandinavian glaciers by their great instability during the Eemian or Holocene interglacials. This instability is inherent to the underlying volcanism and also to the climate, due to the presence of two major warm sea currents, the Irminger in the west of the island and the North Atlantic drift (NAD) in the east. It is also marked in glacial context for the same reasons (interstadial warmings), thus limiting the volume, but not the extent, of the ice caps during the LGM (Last Glacial Maximum) from 60 ky ago. Volcanism, exacerbated by the gravitational spreading of the ice caps during warming and the glacio-isostatic unloading, correlates well, on a large scale, with these different phases of deglaciation, from 30 ky in the north of the island. In the south and west of Iceland, the accentuation of volcanism occurs earlier in relation with a more marked exposure to the thermal input of the Irminger current and the consequent emergence from the ice cap of the large central volcanoes such as Hekla or Katla. The frequency of eruptions at the beginning of the Holocene was approximately twice as high as the average for the whole postglacial period, although their frequency increased again at the end of this final period. This phenomenon is attributed to coupled changes in tectonic and magmatic regimes during glacial unloading (Guðmundsson 1986, 2000; Sigvaldason et al. 1992; Pagli and Sigmundsson 2008; Edwards et al. 2015). During the last interglacial, the situation was similar to that of the Holocene. During glaciation, there is in theory a partial confinement of magma at the hotspot level; its release occurs during deglaciation. Nevertheless, Grímsvötn remained regularly active throughout the last glaciation, indicating a continuous deep magma feeding. The beginning of the postglacial period coincided mainly with effusive volcanism and only a few explosive eruptions for the most active Icelandic volcanoes. However, this was not the case for the west volcanic zone (WVZ), fed

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directly by the Atlantic Ridge, and rather in the process of extinction during a full interglacial period (< 8,000 years BP). 2.2.1.2. Plausible explanations for variations in partial melting Numerous studies document the magmatic impact of the glacio-isostatic unloading, linked to the melting of the ice cap and thus to the rapid disappearance of its stress on the lithosphere (section 3.2) at the beginning of interglacial period. This discharge by adiabatic lowering of the applied pressure induced an increase in the partial melting rate of the upper mantle when the latter was decompressed (Jull and McKenzie 1996; Maclennan et al. 2002; Eason et al. 2015). Indeed, in dry conditions or in the presence of small quantities of fluids, the solidus (curve marking the very beginning of melting) of the peridotites and pyroxenites of the upper mantle shows a positive slope in the pressure (ordinate)-temperature (abscissa) diagrams. Therefore, under adiabatic conditions (i.e. at constant temperature), any decompression of the partially melting peridotites tends to move them away from their solidus towards the domain of increasing liquid contents, leading to the appearance of larger quantities of basaltic magmas. In addition, the incompatible elements in mantle rocks are mainly hosted by clinopyroxene, which is the most easily melted mineral in the dry upper mantle. As a result, the incompatible element contents of basaltic liquids will be highest in magmas with low melting rates (less than 5%), and will tend to decrease when this rate increases following decompression. It can thus be seen, for example in Katla (Figure 1.7), that early postglacial lavas were depleted in incompatible elements such as the rare earths compared to those emitted during the glacial period. This increase in the melting rate during a deglaciation episode induces a strain release within the rifts, linked to the injection of magmas in the form of dykes fed by very large subcrustal basaltic reservoirs, and then to their injection into the shallower reservoirs underlying the central volcanoes. It is not necessary to postulate a considerably increased melting rate during glacio-isostatic discharge. Jull and Mackensie (1996) suggest an increase in the average melting rate of only 0.2%. The increase of mantle melting may also be partly controlled by decompression related to surface crustal extension in rift zones (Maclennan et al. 2001). Modeling based on deglaciation modalities similar to those described in this book (Chapter 3) estimates that an uplift rate of the partial melting zone of 100 m/year (Eksinchol et al. 2019) is enough to explain the temporary increase of the eruptions, in relation to the different deglaciation episodes (Figure 2.9), and in particular to explain the Saksunarvatn polyphased eruptive crisis (four tephra ranging in age from 10.3 to 9.8 ky).

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Figure 2.9. Comparative volumes of volcanic emissions during the end of deglaciation and the Holocene. The volume emitted was maximum immediately after the end of the deglaciation of the island, around 8.3 ky (modified from Hjartasson 2003)

In fact, almost all of Icelandic very large volcanic areas are rift-related, including those of Grìmsvötn-Laki and Bárðarbunga-Veiðivötn. A complementary decompression factor other than ice melting is an increased erosion rate during deglaciation. Finally, a last possible discharge factor is directly related to the emptying of crater lakes or the drainage of glacial aquifers. Deglaciation also causes stress release and reactivation of faults (Mörner 1978; Turpeinen et al. 2008) within and beyond the ice cap limits (James and Bent 1994). It is generally accepted (Jakobsson et al. 1978; Sigvaldason et al. 1992; Slater et al. 1998) that periods of increased volcanic activity in Iceland are related to crustal extensional regime, itself partly controlled by glacial discharge and seismicity (Guðmundsson 1986; Hasegawa and Basham 1989; Johnston 1989; Van Vliet-Lanoë et al. 2005). Induced seismicity also occurs within 1,000 years following deglaciation (Hasegawa and Basham 1989). Guðmundsson (1986) also suggested that the variation in crustal stress due to glacio-isostatic unloading increases fracturing of the magma chambers and thus increases eruption rates. Geochemical variations of incompatible elements, less abundant in postglacial lavas than in glacial lavas (Slater et al. 1998), and isotopes (displaying more primitive signatures in postglacial lavas (Gee et al. 1998)) are consistent with such models of

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increased magma production and partial mantle melting rate during deglaciation (Jull and McKenzie 1996; Eason et al. 2015; Edwards et al. 2015). These variations, widely accepted by Icelandic igneous geochemists (Harðarson et al. 2008; Huybers and Langmuir 2009), have been demonstrated in numerous volcanoes, for example Dyngjufjöll (Sigvaldason et al. 1992), Krafla (Slater et al. 1998; Maclennan et al. 2002), Grímsvötn (Albino et al. 2010), Katla (Sims et al. 2013) (Figure 1.7) and the whole West Volcanic Zone (Sinton et al. 2005; Eason et al. 2015). During deglaciation, a major basaltic eruption can durably modify the subsequent rate of magmatic production; for example, the Katla eruption rate decreased from an average of 2 km3/century before the Eldgjá eruption in 939 AD to only 0.7 km3/century since this date (Oladottir et al. 2018). 2.2.2. Deglaciation and climate feedback The impact of volcanism on climate changes and the effect of climate changes on volcanism are now well documented, but remain a subject of discussion (Cooper et al. 2018; Harning et al. 2018; Swindle et al. 2018a, 2018b). It is well established that major volcanic eruptions have an impact on the climate both in interglacial and glacial periods. Their impact is primarily related to the emission of gases such as CO2 and especially SO2 which is transformed into sulfuric aerosols. These aerosols and the load of volcanic ash in the upper atmosphere decrease the energy absorption in the troposphere and at ocean and ground level. An emblematic Icelandic example is the fissural eruption of Laki in 1783–1784, which released huge amounts of sulfur-rich gases into the stratosphere. On the other hand, the increase in CO2 production linked to subaerial volcanism seems sufficiently significant to influence glacial/interglacial CO2 variations and thus climate transitions (Huybers and Langmuir 2009). The intensification of volcanic activity during deglaciation results, as discussed above, from readjustments of the crustal tension associated with (1) the post-glacial redistribution of the lithosphere masses, but also (2) that of the masses of marine waters, under control of terrestrial obliquity and eustatism. The increase of volcanism in the northern hemisphere was delayed by 3 ky compared to the acceleration of the global eustatic rebound at the beginning of the Holocene (Melt Water Pulse 1b). This is in relation to orbital forcing, under the primary control of the warming of the southern hemisphere (Dansgaard–Oeschger events) ((Van Kreveld et al. 2000; Rasmussen et al. 2016); see Chapter 3). Volcanic eruptions can also have a direct retroactive impact on the mass balance of ice caps due to induced climatic changes (loss of altitude, modification of

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atmospheric circulation). They also impact the radiative emissions from the ice, due to the loss in albedo caused by the deposition of fine-grained volcanic ashes at high latitudes (pack ice, ice caps) (Praetorius et al. 2016). Analysis of deglaciation varves shows that abrupt ice cap melting events coincide with volcanic aerosol emissions, such as those recorded in Greenland ice cores (Zielinksi 2000; Kutteroff et al. 2013: Gettelman et al. 2015). 2.3. The hypothesis of a link between volcanism and climate and its test by dating During the last two decades, geochronological studies have contributed to testing the hypothesis of a possible link between volcanism and deglaciation (McGarvie et al. 2006, 2007; Flude et al. 2008, 2010) and refining our knowledge of the extent of the ice sheet during the Icelandic glaciations (Hoppe 1982; Einarsson et al. 1988; Geirsdottir et al. 1997; Ingólfsson et al. 1997). To test the link between deglaciation and volcanism over periods extending beyond the last 20,000 years, it is necessary to date a representative set of eruptions in subaerial and subglacial contexts, respectively. Two of the dating methods used have only been able to investigate the last 20,000 years of Icelandic volcanic activity. These are the radiocarbon (14C) method applied to the dating of charcoal buried under lava flows (Sinton et al. 2005) (see Box 2.1) and the method based on concentrations of a stable cosmogenic isotope of helium, 3He (Licciardi et al. 2007). The latter measures the duration of exposure of the lava surface to cosmic bombardment. In a simplified way, this duration is proportional to the 3He content of the crystals occurring at the exposed surface of a subaerial lava. Licciardi et al. (2007) considered that this exposure time was equivalent to the age of the last eruptions of the subaerial lavas capping the tuyas. Thus, some of the most recent subglacial eruptions in Iceland could be dated. One of the limitations of this method is that Iceland was repeatedly recovered by glaciers or snow packs that form a shield against cosmic bombardment. Thus, a 3He age dates the last glacial melting event, that is the re-exposure of the lava to the cosmic bombardment, rather than the crystallization of the lava on the surface. The 14 C method is limited to the dating of eruptions younger than 50,000 years, considering the relatively short half-life (5,734 ± 40 years) of this isotope (see Box 2.1). It also requires the occurrence of carbonized material at the base of the flows, which may be a limiting factor, as wood was scarce in Iceland except during the interglacial optima. Methods for dating older periods rely mainly on the

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K-Ar chronometer, whose success rate is limited by the potassium depletion of most Icelandic lavas. Carbon-14 dating is a radiometric dating method based on the measurement of carbon-14 (14C) contained in organic matter to determine its absolute age, that is the time elapsed since its death. The method’s field of use corresponds to absolute ages from a few hundred years to, at most, 43,000 years. Carbon-14 is one of the radioactive isotopes of carbon with a half-life of 5,734 ± 40 years. For dating, the 1951 estimate of 5,568 ± 30 years is used by convention. The most common method of dating is to determine the radiocarbon concentration (i.e. the ratio 14C/Ctotal) of a sample at time measurement. Natural radiocarbon circulates in three reservoirs: the atmosphere, the oceans and the biosphere. It comes mainly from the very high atmosphere where it results from the decomposition of atmospheric nitrogen under the impact of cosmic radiation. It combines with oxygen to form CO2. It is then absorbed by plants during photosynthesis, incorporated into the food or trophic chains, and then fossilized after the death of the organism. Its quantity then starts to decrease exponentially according to the radioactive decay process, which makes it possible to calculate how long the organism has been dead. It also dissolves in the oceans and co-precipitates to form carbonates. The 14C/Ctotal ratio is considered to be uniform in the atmosphere, the ocean surface and the biosphere because of the permanent exchanges between living organisms and their environment. The integration of atmospheric carbon by marine waters takes time and the global ocean circulation is a long cycle (about 1.5 ky), which includes deep currents whose upwelling brings to the surface waters several hundred years old. These phenomena imply that the carbon in ocean waters is on average 400 years older than that in the atmosphere (global reservoir effect). Box 2.1. Carbon-14 or radiocarbon dating

2.3.1. The K-Ar chronometer 2.3.1.1. Principle Two reference works (Dalrymple and Lanphere 1969; McDougall and Harrison 1988) detail the operation and conditions of use of the K-Ar and 40Ar/39Ar methods. The principles necessary for their application to the dating of Icelandic lavas and tephras are therefore presented here. One of the isotopes of potassium, 40K, which represents only 0.01167% of total potassium, is naturally radioactive. It has a half-life of 1.25 × 109 years. About 89%

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of 40K decays to 40Ca, 11% to 40Ar* (* means here radiogenic, i.e. from the radioactive decay of 40K). The decay of 40K (parent isotope) into 40Ar* (daughter isotope) is the basis of the chronometer. At high temperatures (1,200°C for a basaltic magma), the 40Ar* gas is continuously evacuated from the magma. Rapidly after eruption, whether effusive – producing lava flows – or explosive – resulting in the ejection then deposition of tephra – the magma solidifies. 40K inexorably continues its disintegration into 40Ar*, but the 40Ar* will remain trapped in the crystal lattice after the solidification of the magma. Thus, by jointly measuring the number of remaining parent isotopes (40K) and the number of daughter isotopes formed (40Ar*) and knowing the decay constants of 40K, we can calculate the time elapsed since the date of the eruption, in other words the age of the rock. This method implies the following assumptions. The rock or crystal is considered to evolve within a closed system, from solidification to dating. This means that the sample must not have been reheated, which would result in an unquantifiable loss of 40Ar*, or altered, which would also result in a loss of 40Ar* accompanied with a loss or gain of 40K depending on the type of alteration (Guillou et al. 2017). In a context such as that of Icelandic volcanism, the interaction between ice and magma, hydrothermal alteration, meteoric alteration are all external factors promoting the loss and/or gain of one or both isotopes and thus the disruption of the K-Ar clock. It is therefore necessary to ensure the absence of alteration by means of macroscopic and microscopic observations and geochemical analyses, including in particular the water content, which provides valuable information about the “freshness” of the samples. It is also postulated that at t = 0 (date of eruption and crystallization of the lava), the magma is devoid of 40Ar*. Thus, the potential argon initially trapped in the sample is considered to have a 40Ar/36Ar ratio equivalent to that of the atmosphere, that is 298.56 (Lee et al. 2006). As a reminder, 36Ar is exclusively of atmospheric origin, as no naturally occurring radioactive isotope produces 36Ar as a decay product. Therefore, a K-Ar age will be overestimated if, at t = 0, that is at the time of eruption, the sample already contains 40Ar* not resulting from the disintegration of the 40K groundmass that crystallized after the closure of the system. Thus, at t = 0, the 40Ar/36Ar ratio will be greater than 298.56, the value assigned to the atmospheric argon ratio. This argon, commonly called “foreign argon” (Dalrymple and Lanphere 1969), is of two types. During an eruption, xenoliths of ancient rocks from the surrounding bedrock, or even foreign crystals (xenocrysts), can be incorporated into the magma as it rises to the surface. These xenoliths and xenocrysts are necessarily

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older than the eruptions that brought them to the surface. This argon is referred to as inherited argon. Sample preparation protocols are designed to eliminate, or at least to minimize this source of error. The removal of these crystals or lithic fragments from the dated fraction is done by densitometric and magnetic sorting, to retain only the groundmass, which is the part of the lava crystallized during its cooling at surface. The second type of argon is “excess” argon. It is the fraction of argon 40 introduced into the sample during its crystallization: it does not come from the in situ disintegration of the 40K within the sample. 2.3.1.2. Experimental approaches The origin of this excess argon is still poorly understood. It is generally trapped within fluid or glassy inclusions in the peripheral zones of the constituent grains of the rock (Kelley 2002) and it is not degassed from the magma when the latter cools at the surface. Dating of recent Hawaiian submarine lavas has shown that glassy samples should also be avoided, due to the possible excess argon caused by rapid quenching combined with too high hydrostatic pressure preventing the loss of trapped argon that may not have the composition of atmospheric argon (Dalrymple et al. 1999). Thus, the presence of excess argon is possible in lavas constituting móbergs (flat-topped volcanoes built up under a thick layer of ice) and hyaloclastite ridges. Densitometric and magnetic sorting are ineffective in separating this argon from the purely radiogenic argon formed in situ. The sources of error can however be reduced by not retaining the glassy products which, in addition to their tendency to present excess argon, are easily altered, and by working only on the internal and crystallized parts of the subaqueous lavas. A complete and detailed description of the two methods based on the K-Ar chronometer is available in Duplessy and Ramstein (2013). Here is an outline of the two methods. 2.3.1.3. Sample preparation Only unaltered samples will be retained, in particular to satisfy the closed system evolution hypothesis discussed above. For example, a basalt is considered unaltered if its water content is less than 1%. The dated part (mineral, glass, etc.) of the selected sample must be representative of the event to be dated. In the case of volcanic rocks, the phase selected will be for lavas their microcrystalline groundmass, because it formed by rapid cooling when the magma reached the surface; for tephras, the phenocrysts (feldspars, micas or amphiboles) synchronous with the eruptive event (i.e. juvenile) are the only ones retained. After field sampling, the rocks are sawn, crushed and

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sieved to the 0.250–0.125 mm fraction size. The resulting aggregate is ultrasonically washed in a bath of acetic acid (1N) heated to 50°C. The sample is then rinsed thoroughly with deionized water. Separation of groundmass or phenocrysts is done in the laboratory by magnetic and densitometric separation or manual sorting under a binocular magnifying glass. In the case of tephras, the preparation method is identical. Given the nature of the deposit, however, the grinding phase is not necessary. 2.3.1.4. The unspiked K-Ar method The K-Ar dating method used to date the Icelandic lava samples was the unspiked K-Ar technique developed by C. Cassignol (Cassignol et al. 1978, 1982). To calculate an age, the measurement of only two variables is necessary: the radiogenic argon 40 content (40Ar*) and the potassium 40 content (40K). Potassium is a major element whose analysis is done by inductively coupled plasma atomic emission spectroscopy (ICP-AES). This is a classical method. The analysis of the isotopic composition of argon is preceded by the extraction of gas from the sample, then its purification. The sample (0.5 to 3 g depending on the case) is dropped into a molybdenum crucible placed in an ultra-vacuum chamber. The sample is melted under vacuum using an induction furnace. Then the extracted gas is introduced into an ultra-vacuum line connected to a mass spectrometer. This gas is then purified using a Titanium sublimation pump and under the action of getters pumps. The gas thus purified is expanded in the ultra-vacuum line, attracted by condensation on a nearby activated carbon finger, then introduced into the mass spectrometer cell. Argon, which is a neutral gas, is ionized by an electronic source. 40Ar becomes Ar+ and 36Ar becomes 36Ar+. The now charged atoms are accelerated under the effect of a potential difference. They then pass through a magnetic field following a circular trajectory. In a high vacuum chamber, these ions of mass 40 and 36 and charge, animated by a velocity acquired by a potential difference V, describe, during their passage through a magnetic field H, a trajectory whose radius is inversely proportional to the mass of the isotope. 40Ar and 36Ar isotopes of the sample (Arech) are measured simultaneously on a double collector composed of two Faraday collectors arranged at m/e = 40 and m/e = 36. The signals are integrated over a period of 100 seconds. Once the gas analysis is complete, the sample is evacuated from the mass spectrometer by cryo-pumping. The reference atmospheric argon (Arat) is introduced from a cylinder containing laboratory air (atmospheric reference), then measured under the same conditions pressure as the sample, allowing direct comparison of the two gas aliquots (sample and atmospheric reference) and determination of the relative radiogenic argon content (Figure 2.10). 40

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The 40Ar* content is given by the equation: 40

Arech 40 Arat − 36 Arech 36 Ar at 40 * Ar = 40 Arech 36 Arech

Figure 2.10. Principle of the unspiked K-Ar method. Two isotopes of argon are measured with the mass spectrometer (40Ar and 36Ar). The sample gas is composed of 40Arat and 36Arat of atmospheric origin, as well as 40Ar* radiogenic resulting from the 40K radioactive decay; cd: calibrated dose

The third measurement of an aliquot of gas is done for calibration, that is the conversion of an electrical signal into a number of atoms. A known number of 40Ar atoms are introduced from a calibrated cylinder into the mass spectrometer. The number of atoms (cd) is deduced from the measurement of standard minerals of known age. The procedure for establishing the calibration curve equation is detailed in (Charbit et al. 1998). The three previous steps are shown in Figure 2.10. 2.3.1.5. The 40Ar/39Ar method In this method, which is a variant of the K-Ar method, the samples undergo neutron activation before the mass spectrometric measurements. The samples are activated via a flux of fast neutrons within a nuclear reactor, to artificially transform part of the 39K isotopes into 39Ar. The amount of 39Ar thus generated is proportional to

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the number of 39K and thus 40K (parent isotope) present in the sample, the ratio 40 K/39K being (assumed) constant in nature. The advantage is that the ratio between radiogenic daughter atoms (40Ar*) and parent atoms (39Ar proportional to 40K) is obtained by a simultaneous mass spectrometry measurement. 39Ar is radioactive. Its decay period is 265 years. As the spectrometric analyses are carried out within less than one year after irradiation, the error in its determination is negligible. Precise knowledge of the production yield of 39Ar from 39K is obtained by referring to standards of known ages, irradiated together with the samples to be dated. For our studies on Icelandic lava and tephra, these are the sanidine crystals from the Alder Creek rhyolite (USA), dated at 1.193 ± 0.001 My by Nomade et al. (2005). The advantage of producing 39Ar at the expense of the parent element (40K) is that this transformation makes it possible to substitute the direct measurement of the ratio 40 Ar/39Ar (by mass spectrometry) to that of the ratio 40K/40Ar (by two different methods, atomic absorption for 40K and mass spectrometry for 40Ar). 2.3.1.6. Step-heating experiments This type of experiment is used for the dating of lavas. After irradiation, the microcrystalline matrix of the sample (groundmass) is placed on an under-vacuum sample holder connected to a furnace or a laser, themselves connected via an ultrahigh vacuum line to a mass spectrometer. The sample is gradually heated in increasing temperature steps (e.g. 60°C steps). For each step, the argon isotopic composition of the extracted and purified gas is measured with the mass spectrometer. An apparent age can thus be calculated for each step of sample degassing. The final result is an age spectrum (Figure 2.11). The general appearance of this spectrum shows whether or not the K/Ar clock within the sample has been disturbed. An undisturbed sample that has evolved in a closed system since its crystallization will present at each temperature step a constant ratio between the 40 Ar* and 39Ar isotopes. We will thus obtain, for each temperature step, identical apparent 40Ar/39Ar ages within analytical errors. The result of the experiment will be represented as an age spectrum where each temperature step will give the same age. The age of the sample can then be defined as the plateau age. It is generally considered that a plateau is composed of at least three successive steps containing at least 60% of the degassed 39Ar and whose apparent ages are consistent at 95% probability (Sharp and Renne 2005). Alteration, metamorphism and hydrothermalism are at the origin of migrations, losses and/or gains of argon and potassium isotopes and will be highlighted by disturbed age spectra that show non reproducible ages from one step to another.

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For samples containing argon in excess, the 40Ar/39Ar step-heating experiments evidence the degassing of fluid inclusions at low temperatures. They yield ages that are too old in the early steps of the age spectrum. At high temperatures, glassy or mineral inclusions also give ages that are too old in the later steps of the age spectrum (Esser et al. 1997). The discordant, U-shaped age spectrum is then symptomatic of excess 40Ar. Fluid inclusions and glassy inclusions are frequent sources of excess argon, and can significantly impact the dating of rocks, especially those that are young and/or low in potassium. The argon richness of these inclusions can be explained by the incompatible behavior (i.e. the strong affinity for fluid or liquid phases) of argon. Thus, glassy inclusions originating from a magma rich in excess argon can contain up to 100 times more argon (by mass) than minerals crystallizing from the same magma. In the case of fluid inclusions, this ratio can reach up to 10,000 (Kelley 2002).

Figure 2.11. Examples of age spectra and isochron diagram

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COMMENT ON FIGURE 2.11. – (A) Undisturbed age spectrum. Each rectangle corresponds to a temperature step. All steps are retained in the calculation of the plateau age. On the x-axis, the 39Ar released is an indicator of the degassing of the sample, ranging from 0% at the beginning of melting to 100% when the sample is fully melted. (B) Disturbed spectrum. The first two steps show excess ages probably due to excess argon and/or loss of K by weathering. Younger ages at high temperatures indicate loss of argon. (C) Inverse isochron diagram. All the steps are retained, the isochron age is equal to the plateau age. The 40Ar/36Ar value at the intercept is 299.7 ± 2.5, a value equivalent to the atmospheric composition. There is therefore no excess argon for this sample. The 40Ar/39Ar method can also be used to process the isotopic data obtained to construct isochrons. The experimental values 39Ar/40Ar versus 36Ar/40Ar acquired for each temperature level are plotted in a diagram called “inverse isochron”. When the extracted gas is a simple atmospheric argon-radiogenic argon mixture, the points are aligned along a mixing line or isochron. The intersection of the isochron with the 36 Ar/40Ar axis corresponds to a value of 39Ar equal to 0. This deduced and calculated value corresponds to the isotopic signal of a sample devoid of K and which cannot therefore produce purely radiogenic 40Ar*. Thus, for a sample devoid of excess 40 Ar*, the isochron intercepts the 36Ar/40Ar axis at an ordinate of value 1/298.56 (Lee et al. 2006), which we will note as (36Ar/40Ar)i.. Conversely, for a sample containing excess 40Ar*, (36Ar/40Ar)i will have a value of less than 1/298.56. Thus, the presence or absence of excess argon can be seen from the inverse isochron diagram. 2.3.1.7. Single grain dating experiments This type of experiment is implemented for the dating of tephra and other products of explosive volcanism. After irradiation, the selected phenocrysts, generally feldspars, are placed under vacuum on a sample holder. They are then melted individually or in groups of three or four identical and homogeneous crystals (we then speak of population), using a laser. The extracted gas is then purified by getters pumps and analyzed by mass spectrometry. An age is obtained for each crystal or for each constituent population of the tephra level. Thus, age probability diagrams can be established for a given tephra (Deino et al. 1992). Analysis of these diagrams allows the homogeneity of the stratigraphic level to be estimated and the most statistically probable age to be defined (Figure 2.12).

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Figure 2.12. Example of probability diagrams (continuous red curve) of anorthoclase crystals and their individual 40Ar/39Ar age distribution

COMMENT ON FIGURE 2.12. – The gray curve is drawn taking into account all the experiments. It shows several populations of grains of different ages. The red curve is drawn by retaining only the youngest ages which define a unimodal Gaussian distribution. N = number of crystals retained in the age calculation compared to the number of crystals analyzed.

Figure 2.13. Isochron diagram showing in gray the xenocrysts that plot off the isochron (black), unlike the juvenile crystals from the erupted magma

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The analysis of the results by the isochron approach is particularly useful for the dating of tephras. Indeed, in an ideal case, all minerals of the same tephra level should plot on the same isochron, because they have a priori the same age. If some experimental points do not plot on this isochron, we can deduce that the corresponding crystals are xenocrysts. This means that they are either older crystals or crystals of foreign origin, remobilized during the eruptive event at the origin of the tephra (Figure 2.13). 2.3.2. The combination of K-Ar and Icelandic volcanism

40

Ar/39Ar methods for dating

2.3.2.1. Methodological considerations 40

Ar/39Ar dating is more complex to implement and achieve than K-Ar dating. Indeed, it requires neutron activation of the samples, involving a set of corrections related to this irradiation that will not be discussed here but is detailed in (McDougall and Harrison 1988), including joint analysis of minerals to calibrate the neutron flux and tedious step heating experiments. For this reason, it is preferable to begin the chronological study of a volcanic region by using the K-Ar method, which is much quicker to implement and allows for establishing a first chronological reference frame. This approach was followed by Guillou et al. (2010, 2019) to date recent Icelandic volcanism. The application of the K-Ar chronometer to the dating of Icelandic volcanism poses some difficulties that can be circumvented by using the ad hoc methodological approach(es). First, the great majority of this volcanism is of tholeitic affinity. Such lavas are potassium-poor and thus have low contents of the parent isotope (40K). Moreover, for a recent volcanism (< 500 ky), the time between the subaerial cooling of the magma and the dating of the sample is short at the geological scale. The chronometer is therefore started with a low number of parent isotopes, and the duration of radioactive decay is relatively short. The combination of these two factors means that the number of radiogenic atoms formed will be small. The dating technique to be used must therefore ensure a fair and accurate measurement of low levels of the radiogenic isotope (40Ar*). The unspiked K-Ar dating technique as described in (Charbit et al. 1998) meets this condition. This makes it a tool of choice for the dating of Quaternary effusive subaerial volcanism. However, this method does not allow the separation of the 40Ar* resulting from the disintegration of 40K from the groundmass of the excess 40Ar (Sasco et al. 2017). It also does not allow us to know if the system remained closed with respect to 40K and 40Ar* over the time elapsed from surface cooling to laboratory analysis of argon and potassium isotopes.

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In unspiked K-Ar dating, the entire extracted argon is analyzed by mass spectrometry in a single shot. Unlike the 40Ar/39Ar method, K-Ar dating does not allow for the possibility of highlighting possible re-openings of the isotopic system and of distinguishing radiogenic argon from excess argon by means of isochrons and age spectra. Therefore, K-Ar ages should always be viewed with caution and we should accept the fact that the actual error on such ages can be much greater than analytical error. Nevertheless, studies comparing 40Ar/39Ar and K-Ar have shown that for fresh samples without excess argon, the unspiked K-Ar method gave results consistent with those of the 40Ar/39Ar method (Singer et al. 2009; Guillou et al. 2011, 2016; Laj et al. 2014). 2.3.2.2. Sample selection In order to test the possible link between volcanism and climate, samples from subglacial eruptions and subaerial eruptions must be dated (Figure 2.14). Due to the relatively large number of samples concerned, the first dating method applied is K-Ar. Subglacial samples are taken from two types of subglacial volcanic bodies: table mountains, also known as tuyas or móbergs, and hyaloclastite ridges (see Chapter 1 and section 2.1; Figure 2.15). Detailed descriptions of the lithofacies and sequence structure of subglacial volcanoes are available in numerous publications, for example (Werner and Schmincke 1999; Helgason et al. 2001; Komatsu et al. 2007), and summarized in (Smellie 2008). Tuyas are the subglacial equivalent of subaerial shield volcanoes (Werner et al. 1996). Hyaloclastite ridges are the subglacial equivalent of fissural volcanoes and are aligned along the main tectonic directions. During warmer periods, subaerial eruptions dominate. Compared to subglacial eruptions that have produced high topographic bodies of limited lateral extent, subaerial eruptions tend to cover and smooth the topography. During these eruptions, very large volumes of lava from fissures or eruptive shields can be emplaced on the surface. Their volumes can occasionally exceed 20 km3 (Lakí lava flow field). Dated subaerial samples (Guillou et al. 2010) come from pahoehoe and aa type lava flows. The columnar-jointed basalts of Gerðuberg are a spectacular example of subaerial volcanic activity (Figure 2.15(E)).

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Figure 2.14. Location map of dated samples (modified from Guillou et al. 2010)

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Figure 2.15. Examples of dated subglacial lava flows (A–D) and subaerial lava flows (E–F) (© Hervé Guillou)

COMMENT ON FIGURE 2.15. – (A) ISLN-65, Þórósfell tuya, details of the jointed part of the hyaloclastite unit. (B) ISLN-43, intrusion into the hyaloclastite formation of the Vatnafell tuya. (C) ISLN-59, massive rounded block of the hyaloclastite crest, Upptyppingar. (D) ISLN 64, jointed flow from the Laufafell tuya. (E) ISLN-46, columnar-jointed Gerðuberg subaerial lava flow. (F) ISLN-56, upper part of a partially dismantled subaerial lava flow near Mulalón (from Guillou et al. 2010).

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2.3.2.3. Results of K-Ar dating Sample nb.

40

Ar* ± 1 (10-13 σ weighted (%) mol/g) average 0.103 1.21996 0.581 ± – – – 2.88534 0.006 0.040 – 1.24136 0.266 ± 0.008 – – 1.21005 0.003 – 0.209

Melted Measurement K* weight nb. (wt. %) (g)

40Ar*

40Ar*

Age ±2σ ky

ISLN-65: Þórósfell tuya

7.610 7.626

ISLN-63: Herðubreið tuya

7.609 7.693

ISLN-59: Upptyppingár ridge ISLN-17 Skarðsengi basalt on fm Sydra ISLN-24 Sandfell tuya ISLN-22 Kárahnjúkur: (tunnel): 525 m asl. SN-10 Snaefell peak lava (1,533 ml) ISLN-60A: Herðubreiðartögl tuya ISLN-64: Laufafell tuya

7.440 7.694

1.00685 0.581 ± 0.568 0.486 2.49684 0.006 0.463 0.477

0.478 ± 0.036

48 ± 7

6.157 6.181

1.49664 0.522 ± 0.314 0.762 1.06354 0.005 0.337 0.676

0.731 ± 0.041

81 ± 9

6.887 6.903

1.79824 0.631 ± 0.414 1.674 1.72547 0.006 0.391 1.616

1.638 ± 0.046

150 ± 9

6.844 6.868

1.56899 0.589 ± 0.531 2.367 1.14643 0.006 0.544 2.673

2.459 ± 0.057

241 ± 12

7.196 7.211

1.07774 3.686 ± 3.665 16.951 1.45909 0.037 3.271 15.805

16.193 ± 0.092

253 ± 6

7.637 7.712

1.50960 0.133 ± 0.568 0.627 1.08739 0.001 0.757 0.553

0.591 ± 0.075

256 ± 66

7.416 7.426

1.51683 3.130 ± 1.154 20.995 0.98591 0.031 1.190 19.431

20.229 ± 0.166

373 ± 10

ISLN-43: Vatnafell tuya

7.219 7.233

1.22993 0.639 ± 5.922 4.722 2.28773 0.006 4.293 4.535

4.583 ± 0.044

414 ± 11

6.923 6.957

1.06213 0.365 ± 1.367 4.795 1.06165 0.004 1.686 4.481

4.684 ± 0.073

740 ± 27

6.904 6.920

2.68551 0.241 ± 7.029 4.181 2.95490 0.002 6.331 4.237

4.209 ± 0.030

1 007 ± 25

ISLN-21 Ytrikárahnjúkur Base of the recent ridge ISLN-31 Soðarkrokúr

–

–

Table 2.1. K-Ar ages of subglacial volcanic samples (LSCE, France). Age calculations are based on the decay constants of Steiger and Jäger (1977). Data from (Guillou et al. 2010)

Volcanism and Glaciations: Forcings and Chronometers

Sample nb.

Melted Measurement weight nb. (g)

K (wt. %)

40Ar*

(%)

85

40Ar*

(10-13 mol/g)

±1 σ weighted average

Age ± 2 σ ky

40Ar*

ISLN-42

7.547 7.662

1.67964 2.55167

0.897 ± 0.009

1.471 2.029

1.477 1.425

1.440 ± 0.034

93 ± 5

ISLN-46 Gerðuberg lava flow

7.640 7.656

1.51831 2.08575

0.980 ± 0.010

1.923 2.644

2.364 2.246

2.289 ± 0.038

135 ± 5

ISLN-54 lava plateau

7.465 7.481

1.01165 1.63586

0.166 ± 0.002

0.477 0.775

0.395 0.478

0.448 ± 0.521

155 ± 36

ISLN-25 Höfoi lava flow

6.842 6.874

1.59596 2.88896

0.141 ± 0.001

0.843 0.891

0.438 0.424

0.429 ± 0.034

175 ± 28

ISLN-35 Snæfell basal lava flow

7.127 7.143

1.18148 1.69880

1.112 ± 0.011

2.315 3.583

4.927 4.927

4.927 ± 0.062

255 ± 8

ISLN-36 Snæfell basal lava flow

7.128 7.144

1.15370 1.53762

1.187 ± 0.012

3.585 4.481

5.520 5.736

5.649 ± 0.062

274 ± 8

SN-02 Snæfell basal lava flow

7.728 7.804

1.41311 1.03234

0.963 ± 0.010

4.222 5.485

5.361 5.528

5.405 ± 0.062

324 ± 12

ISLN-29 Drafalstadir lava flow

6.859 6.875

1.96951 1.83537

0.183 ± 0.002

1.226 0.988

1.771 1.676

1.732 ± 0.089

546 ± 31

ISLN-55 lava flow at the top of the Múlalón paleolake

7.505 7.521

3.01603 3.00699

0.174 ± 0.002

1.052 1.161

1.708 1.784

1.748 ± 0.028

579 ± 22

ISLN-56 internal lava flow of the Múlalón paleolake

7.474 7.498

1.23746 1.31013

0.407 ± 0.004

4.102 3.304

4.829 4.841

4.836 ± 0.059

685 ± 22

ISLN-58 basal flow of the Múlalón paleolake

7.497 7.506

1.88550 1.97718

0.208 ± 0.002

2.532 2.777

2.671 2.816

2.738 ± 0.039

759 ± 27

Table 2.2. K-Ar ages of subaerial volcanic samples. Data from (Guillou et al. 2010)

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These results confirm the ability of the unspiked K-Ar method to produce very precise ages. This precision is between 10 and 15% for the two youngest samples (ISLN-17 and ISLN-59). Age accuracies for samples dated between 150 and 400 ky and with K contents between 0.5 and 1% range from 2.5 to 6%. Precision as good as 5% is obtained for samples aged 100 ky with a K content of approximately 1% (ISLN-46 and ISLN-42). The ages of subglacial samples range from 1,007 ± 25 ky (sample ISLN-31) to 48 ± 7 ky (sample ISLN-59). Two samples, ISLN-63 and ISLN-65, from the Herðubreið tuya and the Þórósfell tuya respectively, gave a “zero” age. The subaerial lavas are dated between 759 ± 27 ky (ISLN-58) and 93 ± 5 ky (ISLN-42). This extended age range thus allows us to test the volcanic-climate relationship over a significant time span. 2.3.3. A link between volcanism and climate according to K-Ar ages? Lisiecki and Raymo (2005) produced, based on a compilation of δ18O measurements on benthic foraminifera, a reference curve of the glacial and interglacial periods. The results of K-Ar dating are plotted on this curve (Figure 2.16). This allows a direct comparison of the eruption ages with a global paleoclimatic record. The fluctuations of ice cover in Iceland are representative of the variation in global ice volume based on δ18O, as illustrated by Lisiecki and Raymo’s records. Nevertheless, at the present time, 11% of Iceland is still covered by glaciers, and therefore subglacial volcanism can develop during interglacials. Thus, subglacial volcanic products may have radiometric ages corresponding to both glacial and interglacial periods. More likely, subaerial volcanics should have radiometric ages consistent with interglacials (warm periods). Once the ISLN-60A, ISLN-54, ISLN-25 samples whose ages are too imprecise (error greater than the duration of an isotopic stage) have been excluded from the interpretation, we can observe from Table 2.1 that: – three of the subglacial samples (ISLN-64, ISLN-17 and ISLN-59) were emplaced during the transition from a warm period (interglacial stage) to a cold period (glacial stage), and three samples (ISLN-21, ISLN-24, SN-10) during a glacial maximum. Two others (ISLN-43, ISLN-22) correspond to eruptions that occurred during a true deglaciation period; – five of the nine subaerial lavas (ISLN-58, ISLN-56, ISLN-36, ISLN-25 and ISLN-42) were emitted during an interglacial-glacial transition (cooling); – three other subaerial lavas were emitted at the beginning (ISLN-46) or during (SN-02, ISLN-55) deglaciation; – a single subaerial sample (ISLN-29) is synchronous with a glacial maximum.

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Figure 2.16. Plot of the eruption time intervals based on radiometric ages on the reference curve of Lisiecki and Raymo (2005). In red: subaerial eruptions; in blue: subglacial eruptions; in green: intersection zones between subaerial and subglacial ages. The width of the red and blue areas includes the uncertainties related to age

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It thus appears, on the scale of the millennium, that there is no preferential period of volcanic activity. If it seems less, even infrequent during glacial maxima, it can develop during both glacial and interglacial stages. Thus, despite the good precision obtained for most of our K-Ar ages, they do not allow us to establish a clear correlation between climate and volcanism and to demonstrate a systematic link between deglaciation and enhanced volcanism. 2.3.4. A rhyolitic volcanism synchronous with deglaciations? 2.3.4.1. Contribution of the dating of rhyolites Icelandic rhyolites are the result of the evolution by fractional crystallization of magmas of either basaltic, icelanditic or dacitic type (section 1.2). This process requires a residence time in a closed magma system where this evolution can take place. To simplify, it may be considered that the load of glaciers can exert enough stress to lock the magma system. The magma will thus evolve towards the rhyolitic differentiated end-member during the glacial stages. Then, it will be abruptly emitted at the surface at the beginning of deglaciation when the constraints will no longer be sufficient to maintain the volcanic system closed. Based on this hypothesis, subglacial rhyolites are key materials to test the hypothesis of the link between volcanism and deglaciation. Furthermore, rhyolitic volcanism is of great chronological interest because its products, given their explosive character and low alterability, are recorded as tephras in the climatic archives of North Atlantic marine sediments (Lacasse and Garbe-Schönberg 2001; Wallrabe-Adams and Lackschewitz 2003; Austin et al. 2004; Davies et al. 2012) or Greenland ice cap (Ram et al. 1996; Davies et al. 2010; Abbott et al. 2012) and continental sequences (Van Vliet-Lanoë et al. 2020b). The INTIMATE project (Blockley et al. 2012, 2014; Davies et al. 2012, 2014) has demonstrated that tephras are an independent temporal marker of climatic variables and a relevant tool for climate interarchive correlation for the last 130 ky. Dating the rhyolites of the Tindfjallajökull, Kerlingarfjöll and Torfajökull volcanoes, which are the main sources of rhyolitic tephra identified in marine and glaciological archives, is therefore of interest to provide temporal markers and to evaluate their synchronism with deglaciation episodes. Because of their high potassium content, rhyolites are a material of choice for Ar/39Ar dating, especially in Iceland. In this context, four subglacial rhyolitic eruptions were studied by Guillou et al. (2019). The location of these samples is shown in Figure 2.17. 40

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Figure 2.17. Location map of dated rhyolitic samples. The yellow stars refer to the dated samples, the black triangles to the main volcanic centers. EVZ: East Volcanic Zone; NVZ: North Volcanic Zone; WVZ: West Volcanic Zone; RP: Reykjanes Peninsula; FZ: Fracture Zones

2.3.4.2. Geological context of the dated rhyolites 2.3.4.2.1. The Þórsmörk ignimbrite It was first accepted that the Þórsmörk ignimbrite (ISLN-153; Figure 1.28(C)) was derived from the Tindfjöll volcano (Jørgensen 1980). This hypothesis was then questioned by Grönvold et al. (1995), who highlighted a geochemical similarity with the Torfajökull rhyolites. Recent work combining field studies, geochemistry and 40 Ar/39Ar dating (Moles et al. 2018, 2019) confirms that the source of this ignimbrite is Torfajökull. Analysis by LA-ICP-MS (laser ablation – inductively coupled plasma mass spectrometry) indicates that this ignimbrite is similar to the rhyolitic component (II-RHY-1) of the North Atlantic Ash Zone II (NAAZ II) tephra layer (Tomlinson et al. 2010) recognized in marine cores (Lacasse et al. 1996; Wastegard et al. 2006; Brendryen et al. 2011; Abbott et al. 2018) and in Greenland ice cores (Grönvold et al. 1995; Ram et al. 1996; Zielinski et al. 1997). Tephra II-RHY-1 is also a robust anchor point used to correlate different climate archives in the North Atlantic region (Austin et al. 2004; Abbott et al. 2016). It was found at a depth of

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2,359.45 m in the NGRIP ice core and dated at 55.4 ± 1.2 ky using the GICC05 time scale (Svensson et al. 2008). 2.3.4.2.2. The Torfajökull obsidian ISLN-154 is a porphyritic obsidian lava flow containing approximately 10–20 volume% of anorthoclase phenocrysts. It comes from the southern end of Rauðfossafjöll, one of the subglacial rhyolitic tuyas that form the ring of structures which surrounds the Pleistocene rhyolitic central plateau of the Torfajökull volcano. 2.3.4.2.3. The Fannborg rhyolite A third rhyolite (ISLN-95) comes from the base of Fannborg, one of the tuyas located at the northern periphery of Kerlingarfjöll. Grönvold (1972), Stevenson (2004) and Flude et al. (2010) have identified at least 21 different rhyolitic eruptive units at Kerlingarfjöll. The corresponding 40Ar/39Ar ages date rhyolitic activity between 345 ± 17 and 68 ± 21 ky (Flude et al. 2010). 2.3.4.2.4. The Snæffel (East Iceland) The fourth subglacial rhyolite (SN-10) comes from the most recent Snæffel stratigraphic unit (Helgason and Duncan 2005). This feldspar-rich rhyolite has a microcrystalline groundmass and occurs as hyaloclastites. From north to south, the Snæffel forms with the Esjufjöll and Öræfajökull volcanoes a 125 km long volcanic line (Figure 2.17). Available 40Ar/39Ar ages locate the volcanic activity of the Snæffel between 466 and 191 ky. 40Ar/39Ar ages (Helgason and Duncan 2005) date the summit of the volcano around 191 ± 55 ky (plateau age) and 157 ± 90 ky (isochron age), while age K-Ar dates the same unit at 253 ± 6 ky (Guillou et al. 2010). This inconsistency between published ages motivated our interest for new 40 Ar/39Ar datings. 2.3.4.3. Results of the new 40Ar/39Ar datings 2.3.4.3.1. 40Ar/39Ar age of the Þórsmörk ignimbrite A total of 28 populations of 3 to 4 crystals were analyzed by (Guillou et al. 2019). Individual 40Ar/39Ar ages range between 48.8 ± 4.2 and 76.2 ± 3.3 ky. This scatter reflects mixed ages due to loss of argon for the younger ones, foreign argon and/or the presence of xenocrysts for the older ones. The best probability of adjustment coincides with a Gaussian distribution centered on 56.2 ± 1.2 ky, which is the weighted mean value of 21 individual experiments. This age is robust as it is unaffected by excess argon. Indeed, for this sample devoid of excess 40Ar*, the isochron intercepts the 36Ar/40Ar axis at a value equivalent to 1/298.56 (Lee et al. 2006), which is that of atmospheric argon.

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2.3.4.3.2. 40Ar/39Ar age of the Rauðfossafjöll (Torfajökull volcano) The 20 groups of 3 to 4 crystals analyzed define an age range between 70.8 ± 3.2 and 178.9 ± 2.7 ky. The statistically most accurate age is obtained after eliminating the six oldest ages. After removing these older crystal populations, the calculated age is 76.1 ± 1.8 ky. The inverse isochron age (77.0 ± 3.0 ky), calculated from the same population, is equivalent to the weighted average of the individual 40 Ar/39Ar ages. This feature confirms the absence of excess 40Ar* in this population, as shown by the atmospheric composition of the initial argon. The oldest ages are here due to the incorporation of foreign argon and/or the loss of K during secondary (i.e. post-magmatic) alterations. 2.3.4.3.3. 40Ar/39Ar age of the Fannborg (Kerlingarfjöll volcano) The three separate step-heating experiments performed on the ISLN-95 sample show similar patterns, with the first two low-temperature stages giving younger ages than the next ones. These first steps, which show that the isotope system, at least for the low-temperature retention sites, may have been disrupted, are eliminated from the calculation of the plateau ages. There is no sign of an excess of 40Ar*. The retained age, 149.3 ± 3.4 ky, is the weighted average of the three isochron ages. 2.3.4.3.4. 40Ar/39Ar age of the Snæffel summit The first step heating experiment on sample SN-10 indicates potential 40Ar* losses at the lowest and highest temperature retention site (Table 2.3). The alteration effects can therefore be considered minor and the isochronous age of 207.0 ± 10 ky from the weighted average of the three experiments is geologically significant. 2.3.4.4. Relationship between rhyolitic volcanism and deglaciation To compare the possible synchronism between rhyolitic eruptions and de-glaciation episodes, we have plotted in Figure 2.18 the ages of the four rhyolites as well as those previously published (McGarvie et al. 2006, 2007; Flude et al. 2008, 2010; Clay et al. 2015) on Lisiecki and Raymo’s reference curve (2005). Given the relatively large errors on some of the 40Ar/39Ar ages, it is, in some cases, difficult to say whether some rhyolites were emplaced during cold or warm periods. While the relationship between rhyolitic eruptions and deglaciation has been clearly demonstrated for the 11–10.8 ky eruption of Askja S (Sigvaldason 2002; Sigurgeirsson 2016), it does not appear to represent a general pattern. Figure 2.18 shows that rhyolites are emplaced either during glacial stadials or during interstadials and preferably during transitions from warm to cold stages. Indeed, as already

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demonstrated by Flude et al. (2008), subglacial rhyolitic eruptions are frequent in Iceland during interglacial periods and this is also the case for three of the four rhyolites dated by Guillou et al. (2019). Age spectrum Experiment sample nb.

Weight (mg)

Stage used (°C)

39Ar

(%)

Age ± l σ (ky)

Isochron 40Ar/36Ar ±

n/N

lσ intercept

Age ± l σ (ky)

ISLN-95, groundmass FG-1490 to FG-1499

123

728– 1,043

68.4

143.5 ± 3.4

6 of 301.3 ± 4.1 9

FG-1561 to FG-1571

125

795– 1,192

79.9

144.8 ± 3.0

8 of 296.1 ± 1.7 151.1 ± 5.3 10

130

659– 1,195

82.7

145.2 ± 3.7

8 of 296.9 ± 1.3 150.2 ± 5.3 10

144.5 ± 1.9

296.9 ± 2.0 149.3 ± 3.4

FG-1870 to FG-1880 Average

136.0 ± 11.6

SN-10, groundmass FG-1521 to FG-1528

125

722–937

71.3

203.2 ± 5.8

5 of 300.9 ± 4.8 8

182.7 ± 39.8

FG-1594 to FG-1602

125

608– 1,015

100. 0

203.3 ± 5.9

8 of 297.9 ± 1.1 8

210.1 ± 12.0

FG-1851 to FG-1858

125

654–971

97.7

197.2 ± 8.2

7 of 298.1 ± 1.5 8

203.8 ± 24.4

202.0 ± 3.6

298.1 ± 1.7

207.0 ± 10.0

Average

Table 2.3. Table of 40Ar/39Ar data from the temperature step-heating experiments on ISLN-95 and SN-10 samples

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Figure 2.18. Position of the 24 dated Icelandic rhyolites on Lisiecki and Raymo’s reference curve (2005) (from Guillou et al. 2010)

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COMMENT ON FIGURE 2.18. – Error bars are here referred to 1 sigma. MIS refers to marine isotopic stages. Sub-stages 5.1, 5.3, 5.5 are shown on the curve. The red stars are samples dated in Guillou et al. (2019). The orange stars are Torfajökull rhyolites (McGarvie et al. 2006; Clay et al. 2015), the green stars are Kerlingarfjöll rhyolites (Flude et al. 2010), the blue stars are Ljósufjöll rhyolites (Flude et al. 2008) and the black stars correspond to the Prestahnúkur rhyolite (dated at 132 ky by Clay et al. 2015 and at 89 ky by McGarvie et al. 2007). This reinforces the conclusion by Flude et al. (2010) that there is no clear evidence, at least at the latitude of Iceland, that warming is the only trigger for rhyolitic eruptions. Out of the 29 dated rhyolites, only a minority were emplaced during a stadial-interstadial transition. This was the case for one of the Kerlingarfjöll eruptions dated at 247 ± 7 ky (Flude et al. 2010) equivalent to the MIS8/MIS7 transition, as well as the Prestahnúkur rhyolite (Figure 2.3(A)) emplaced during the MIS6/MIS5 transition if the age of 132 ± 9 ky is adopted (Clay et al. 2015). Based on its 40Ar/39Ar age, the Þórsmörk ignimbrite was emplaced at the end of the MIS4/MIS3 transition (see Figure 3.15). This position at the beginning of a cooling phase corresponds to the climato-stratigraphic position of the Þórsmörk ignimbrite ashes observed in the NGRIP drill hole, which intersects Greenland ice sheet. The tephra corresponding to the Þórsmörk ignimbrite is indeed observed at a stratigraphic level corresponding to the transition from GI-15.2 (Greenland interstadial) to GS-15.2 (Greenland stadial NGRIP drill hole). The analysis of the last deglaciation confirms this hypothesis: the rhyolitic tephra recorded in the marine and ice cores as well as the tephra found in European continental peat bogs show a relatively regular frequency of acidic eruptions from 15 ky to 8 ky, without any obvious acceleration (see Figure 1.38). This is also due to the fact that deglaciation is not a rapid process (see Chapter 3), but proceeds by successive heating pulses during Bond events and is therefore relatively gradual. 2.3.4.5. Rhyolitic tephras: temporal markers and interarchive correlation tools Þórsmörk ignimbrite is an important chronostratigraphic marker. The North Atlantic Ash Zone II (NAAZ II) layer, its aerial deposition, is one of the four tephra levels classically selected in stratigraphy for direct correlation between Greenland ice cores and marine sequences (Davies et al. 2012). It is also an important temporal marker as it is synchronous with the transition from GI-15.2 to GS-15.2, a sharp cooling event (Austin et al. 2004) during the last ice age (Blockley et al. 2014). This ignimbrite has also been used to estimate the time lag between the volcanic event

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(the eruption) and the deposition of its products in the North Atlantic basin (Bendryen et al. 2011). In addition, the absolute age of NAAZ II was used as a reference to calibrate the glaciological time scale (Svensson et al. 2008). The production of a very precise radioisotopic age for this ignimbrite is an important contribution to paleoclimatology. The dating of the Fannborg rhyolite (Kerlingarfjöll ISLN-95) provides a new stratigraphic correlation point. A synchronous tephra of marine isotope stage 6 (MIS6) has been identified in the North Atlantic LINK 16 core (Abbott et al. 2014). No age has been assigned to this rhyolitic tephra which is located at 735–737 cm in the LINK 16 core, but stratigraphically it is located during MIS6 prior to the transition to MIS5e, implying an age similar to that of the Fannborg eruption. Tephra LINK 16 and sample ISLN-95 have similar SiO2, TiO2, Al2O3 and CaO values. Although Na2O + K2O values are slightly higher in tephra LINK 16 (8.99%) than in ISLN-95 (7.75%), both samples have comparable Na2O/K2O ratios. Abbott et al. (2014) considered the eruptions of Öræfajökull or Torfajökull as potential sources of LINK 16 rhyolitic tephra. However, in view of the new age determination and geochemical concordances, the Fannborg eruption can also be considered a potential source for the rhyolitic tephra LINK 16. The chemical composition of tephras (Chapter 1) is a basic tool for their correlation, but their dating is the most important marker for stratigraphy. 2.4. References Abbott, P.M., Austin, W.E.N., Davies, S.M., Pearce, N.J.G., Rasmussen, T.L. (2014). Re-evaluation and extension of the Marine Isotope Stage 5 tephrostratigraphy of the Faroe Islands Region: The cryptotephra record. Palaeogeogr. Palaeoclim. Palaeoecol., 409, 153–168. Abbott, P.M., Bourne, A.J., Purcell, C.S., Davies, S.M., Scourse, J.D., Pearce, N.J.G. (2016). Last glacial period cryptotephra deposits in an eastern North Atlantic marine sequence: Exploring linkages to the Greenland ice-cores. Quat. Geochron., 31, 62–76. Abbott, P.M., Griggs, A.J., Bourne, A.J., Chapman, M.R., Davies, S.M. (2018). Tracing marine cryptotephras in the North Atlantic during the last glacial period: Improving the North Atlantic marine tephrostratigraphic framework. Quat. Sci. Rev., 189, 169–186. Ackert, R.P., Singer, B.S., Guillou, H., Kaplan, M.R., Kurz, M.D. (2003). Long-term cosmogenic 3He production rates from 40Ar/39Ar and 40K–40Ar dated Patagonian lava flows at 47°S. Earth Planet. Sci. Lett., 210, 119–136.

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Agranier, A., Maury, R.C., Geoffroy, L., Chauvet, F., Le Gall, B., Aviana, A. (2019). Volcanic record of continental thinning in Baffin Bay margins: Insights from Svartenhuk Halvø Peninsula basalts, West Greenland. Lithos, 334–335, 117–140. Albino, F., Pinel, V., Sigmundsson, F. (2010). Influence of surface load variations on eruption likelihood: Application to two Icelandic subglacial volcanoes, Grímsvötn and Katla. Geophys. J. Int., 181, 1510–1524. Bourgeois, O. (2000). Processus d’extension lithosphérique en Islande. Interactions avec les calottes glaciaires quaternaires. PhD thesis, Université de Rennes I, Rennes. Bourgeois, O., Dauteuil, O., Van Vliet-Lanoë, B. (1998). Subglacial volcanism in Iceland: Tectonic implications. Earth Planet. Sci. Lett., 164(1–2), 165–178. Brendryen, J., Haflidason, H., Sejrup, H.P. (2011). Non-synchronous deposition of North Atlantic Ash Zone II in Greenland ice cores, and North Atlantic and Norwegian Sea sediments: An example of complex glacial-stage tephra transport. J. Quat. Sci., 26, 739–745. Charbit, S., Guillou, H., Turpin, L. (1998). Cross calibration of K-Ar standard minerals using an unspiked Ar measurement technique. Chem. Geol., 150, 147–159. Clay, P.L., Busemann, H., Sherlock, S.C., Barry, T.L., Kelley, S.P., McGarvie, D.W. (2015). 40Ar/39Ar ages and residual volatile contents in degassed subaerial and subglacial glassy volcanic rocks from Iceland. Chem. Geol., 403, 99–110. Cooper, C.L., Swindles, G.T., Savov, I.P., Schmidt, A., Bacon, K.L. (2018). Evaluating the relationship between climate change and volcanism. Earth Sci. Rev., 177, 238–247. Dalrymple, G.B. and Lanphere, M.A. (1969). Potassium-Argon Dating. Principles, Techniques and Applications to Geochronology. W.H. Freeman and Co, San Francisco. Davies, S.M., Wastegård, S., Abbott, P.M., Barbante, C., Bigler, M., Johnsen, S.J., Rasmussen, T.L., Steffensen, J.P., Svensson, A. (2010). Tracing volcanic events in the NGRIP ice-core and synchronising North Atlantic marine records during the last glacial period. Earth Planet. Sci. Lett., 94, 69–79. Davies, S.M., Abbott, P.M., Pearce, N.J.G., Wastegård, S., Blockley, S.P.E. (2012). Integrating the INTIMATE records using tephrochronology: Rising to the challenge. Quat. Sci. Rev., 36, 11–27. Davies, S.M., Abbott, P.M., Meara, R.H., Pearce, N.J.G., Austin, W.E.N., Chapman, M.R., Svensson, A., Bigler, M., Rasmussen, T.L., Rasmussen, S.O., Farmer, E.J. (2014). A North Atlantic tephrostratigraphical framework for 130-60kab2k: New tephra discoveries, marinebased correlations, and future challenges. Quat. Sci. Rev., 106, 101–121. Deino, A. and Potts, R. (1990). Single-crystal 40Ar/39Ar dating of the Olorgesailie Formation, Southern Kenya Rift. J. Geophys. Res., 95, 8453–8470.

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Duplessy, J.-C. and Ramstein, G. (2013). Paléoclimatologie – Tome 1 : trouver, dater et interpréter les indices : enquête sur les climats anciens. EDP Sciences, Paris. Eason, D., Sinton, J.M., Grönvold, K., Kurz, M. (2015). Effects of deglaciation on the petrology and eruptive history of the Western Volcanic Zone, Iceland. Bull. Volcan., 77(6) [Online]. Available at: http://doi.org/10.1007/s00445-015-0916-0. Edwards, B.R., Guðmundsson, M.T., Russell, J.K. (2015). Glaciovolcanism. In Encyclopedia of Volcanoes, 2nd edition, Sigurdsson, H., Houghton, B., McNutt, S., Rymer, H. and Stix, J. (eds). Elsevier, Oxford. Einarsson, T. and Albertsson, K.J. (1988). The glacial history of Iceland during the past three million years. Philos. Trans. R. Soc. Lond., 318, 637–644. Eksinchol, I., Rudge, J.F., Maclennan, J. (2019). Rate of melt ascent beneath Iceland from the magmatic response to deglaciation. Geochem. Geophys. Geosys., 20(6), 2585–2605. Flude, S., Burgess, R., McGarvie, D.W. (2008). Silicic volcanism at Ljósufjöll. Iceland: Insights into evolution and eruptive history from Ar-Ar dating. J. Volcan. Geotherm. Res., 169, 154–175. Flude, S., McGarvie, D.W., Burgess, R., Tindle, A.G. (2010). Rhyolites at Kerlingarfjöll, Iceland: The evolution and lifespan of silicic central volcanoes. Bull. Volcan., 72, 523–538. Gettelman, A., Schmidt, A., Kristjánsson, J.E. (2015). Icelandic volcanic emissions and climate. Nat. Geosci., 8(243) [Online]. Available at: http://doi.org/10.1038/ngeo2376. Guðmundsson, Á. (1986). Mechanical aspects of postglacial volcanism and tectonics of the Reykjanes Peninsula, Southwest Iceland. J. Geophys. Res., 91, 12711–12721. Guðmundsson, Á. (1987). Tectonics of the Thingvellir fissure swarm, SW Iceland. J. Struct. Geol., 9, 61–69. Guðmundsson, Á. (2000). Dynamics of volcanic systems in Iceland. Example of tectonism and volcanism at juxtaposed hot spot and mid-ocean ridge systems. Ann. Rev. Earth Planet., 28, 107–140. Guillou, H., Singer, B., Laj, C., Kissel, C., Scaillet, S., Jicha, B.R. (2004). On the age of the Laschamp geomagnetic event. Earth Planet. Sci. Lett., 227, 331–343. Guillou, H., Van Vliet-Lanoë, B., Guðmundsson, Á., Nomade, S. (2010). New unspiked K-Ar ages of Quaternary sub-glacial and sub-aerial volcanic activity in Iceland. Quat. Geochron., 5(1), 10–19. Guillou, H., Nomade, S., Carracedo, J.C., Kissel, C., Laj, C., Perez Torrado, F.J., Wandres, C. (2011). Effectiveness of combined unspiked K-Ar and 40Ar/39Ar dating methods in the 14C age range. Quat. Geochron., 6(6), 530–538. Guillou, H., Scao, V., Nomade, S. (2016). Blake excursion at Vulcano (Aeolian Islands, Italy): Revised K-Ar and 40Ar/39Ar ages. Quat. Geochron., 35, 77–87.

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Guillou, H., Hémond, C., Singer, B.S., Dyment, J. (2017). Dating young MORB of the Central Indian Ridge (19°S): Unspiked K-Ar technique limitations versus 40Ar/39Ar incremental heating method. Quat. Geochron., 37, 42–54. Guillou, H., Scao, V., Nomade, S., Van Vliet-Lanoë, B., Liorzou, C., Guðmundsson, Á. (2019). 40Ar/ 39Ar dating of the Thorsmork ignimbrite and Icelandic sub-glacial rhyolites. Quat. Sci. Rev., 209, 52–62. Harðarson, B.S., Fitton, J.G., Ellam, R.M., Pringle, M.S. (1997). Rift relocation. A geochemical and geochronological investigation of a palaeo-rift in Northwest Iceland. Earth Planet. Sci. Lett., 153, 181–196. Harning, D.J., Geirsdóttir, A., Thordarson, T., Miller, G.H. (2018). Climatic control on Icelandic volcanic activity during the mid-Holocene. Geology, 46, e443. Hartley, M.E. and Þórðarson Þ. (2013). The 1874–1876 volcano-tectonic episode at Askja, North Iceland: Lateral flow revisited. Geochem. Geophys. Geosys., 14, 2286–2309. Hasegawa, H.S. and Basham, P. (1989). Spatial correlation between seismicity and postglacial rebound in eastern Canada. In Earthquakes at North-Atlantic Passive Margins: Neotectonics and Postglacial Rebound, Gregerson, S. and Basham, P. (eds). Kluwer, Dordecht. Helgason, J. and Duncan, R.A. (2001). Glacial-interglacial history of the Skaftafell region, southeast Iceland, 0–5 Ma. Geology, 29, 179–182. Helgason, J. and Duncan, R.A. (2013). Stratigraphy, 40Ar-39Ar dating and erosional history of Svínafell, SE- Iceland. Jökull, 63, 33–54. Huybers, P.J. and Langmuir, C. (2009). Feedback between deglaciation, volcanism, and atmospheric CO2. Earth Planet. Sc. Lett., 286(3–4), 479–491. Jakobsson, S.P., Jónsson, J., Shido, F. (1978). Petrology of the Western Reykjanes Peninsula, Iceland. Journ. Petrol., 19, 669–705. Jakobsson, S.P., Jonasson, K., Sigurdsson, I.A. (2008). The three igneous rock series of Iceland. Jökull, 58, 117–138. Johnston, A.C. (1989). The effect of large ice sheets on earthquake genesis. In Earthquakes at North-Atlantic Passive Margins: Neotectonics and Postglacial Rebound, Gregerson, S. and Basham, P.W. (eds). Kluwer Academic Publishers, Dordrecht. Jull, M. and McKenzie, D. (1996). The effect of deglaciation on mantle melting beneath Iceland. J. Geophys. Res., 101, 21815–21828. Jørgensen, K.A. (1980). The Thorsmörk ignimbrite: An unusual comenditic pyroclastic flow in southern Iceland. J. Volcan. Geoth. Res., 8, 7–22. Kelley, S. (2002). Excess argon in K-Ar and Ar-Ar geochronology. Chem. Geol., 188(1–2), 1–22.

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Komatsu, G., Arzhannikov, S.G., Arzhannikova, A.V., Ershov, K. (2007). Geomorphology of sub-glacial volcanoes in the Azas plateau, the Tuva republic, Russia. Geomorphol., 88, 312–328. Lacasse, C. and Garbe-Schönberg, C.D. (2001). Explosive silicic volcanism in Iceland and the Jan Mayen area during the last 6 Ma: Sources and timing of major eruptions. J. Volcan. Geoth. Res., 107, 113–147. Lacasse, C., Sigurdsson, H., Carey, S., Paterne, M., Guichard, F. (1996). North Atlantic deep-sea sedimentation of late Quaternary tephra from the Iceland hotspot. Mar. Geol., 129, 207–235. Laj, C., Guillou, H., Kissel, C. (2014). Dynamics of the earth magnetic field in the 10–75 kyr period comprising the Laschamp and Mono Lake excursions: New results from the French Chaîne des Puys in a global perspective. Earth Planet. Sci. Lett., 387, 184–197. Lee, J.Y., Marti, K., Severinghaus, K., Kawamura, K., Yoo, H.S., Lee, J.B., Kim, J.S. (2006). A redetermination of the isotopic abundances of atmospheric Ar. Geochim. Cosmochim. Acta, 70, 4507–4512. Licciardi, J.M., Kurz, M.D., Curtice, J.M. (2007). Glacial and volcanic history of Icelandic table mountains from cosmogenic 3He exposure ages. Quat. Sci. Rev., 26, 1529–1546. Lisiecki, L.E. and Raymo, M.E. (2005). A Pliocene-Pleistocene stack of 57 globally distributed benthic δ18O records. Paleoceanography, 20 [Online]. Available at: http://doi.org/10.1029/ 2004PA001071. Maclennan, J., Jull, M., McKenzie, D., Slater, D., Grönvold, K. (2002). The link between volcanism and deglaciation in Iceland. Geoch. Geophys. Geosys., 3, 1062–1087. McDougall, I. and Harrison, T.M. (1988). Geochronology and Thermochronology by the 40Ar/ 39Ar Method. Oxford University Press, New York. McGarvie, D.W., Burgess, R., Tindle, A.G., Tuffen, H., Stevenson, J.A. (2006). Pleistocene rhyolitic volcanism at the Torfajökull central volcano, Iceland: Eruption ages, glaciovolcanism, and geochemical evolution. Jökull, 56, 57–75. McGarvie, D.W., Stevenson, J.A., Burgess, R., Tuffen, H., Tindle, A.G. (2007). Volcano-ice interactions at Prestahnúkur, Iceland: Rhyolite eruption during the last interglacial-glacial transition. Ann. Glaciol., 45, 38–47. Moles, J.D., McGarvie, D., Stevenson, J.A., Sherlock, S.C. (2018). Geology of Tindfjallajókull volcano, Iceland. J. Maps, 14(2), 22–31. Moles, J.D., McGarvie, D., Stevenson, J.A., Sherlock, S.C., Abbott, P.M., Jenner, F.E., Halton, A.M. (2019). Widespread tephra dispersal and ignimbrite emplacement from a subglacial volcano (TorfaJökull Iceland). Geology, 47, 577–580. Musset, A.E., Ross, J.G., Gibson, I.L. (1980). 40Ar/39Ar dates of eastern Iceland lavas. R. Astron. Soc. Geophys. J., 60, 37–52.

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Nomade, S., Renne, P.R., Vogel, N., Deino, A.L., Sharp, W.D., Becker, T.A., Jaouni, A.R., Mundil, R. (2005). Alder creek sanidine (ACs-2): A quaternary 40Ar/39Ar dating standard tied to the Cobb mountain geomagnetic event. Chem. Geol., 218, 315–338. Oladottir, B., Sigmarsson, O., Larsen, G. (2018). Tephra productivity and eruption flux of the subglacial Katla volcano, Iceland. Bull. Volcan., 80(58) [Online]. Available at: http://doi.org/10.1007/s00445-018-1236-y. Pagli, C. and Sigmundsson, F. (2008). Will present day glacier retreat increase volcanic activity? Stress induced by recent glacier retreat and its effect on magmatism at the Vatnajökull ice cap, Iceland. Geophys. Res. Lett., 35, L09304. Ram, M., Donarummo Jr., J., Sheridan, M. (1996). Volcanic ash from Icelandic ∼57,300 yr BP eruption found in GISP2 (Greenland) ice core. Geophys. Res. Lett., 23, 3167–3169. Russell, J.K., Edwards, B.R., Porritt, L., Ryane, C. (2014). Tuyas: A descriptive genetic classification. Quat. Sci. Rev., 87, 70–81. Sasco, R., Guillou, H., Nomade, S., Scao, V., Maury, R.C., Kissel, C., Wandres, C. (2017). 40Ar/39Ar and unspiked 40K-40Ar dating of upper Pleistocene volcanic activity in the Bas-Vivarais (Ardèche, France). J. Volcan. Geotherm. Res., 341, 301–314. Sharp, W.D. and Renne, P.R. (2005). The 40Ar/39Ar dating of core recovered by the Hawaii scientific drilling project (phase 2), Hilo, Hawaii. Geochem. Geophys. Geosys [Online]. Available at: http://doi.org/10.1029/2004GC000846. Sigvaldason, G.E. (2002). Volcanic and tectonic processes coinciding with glaciation and crustal rebound: An early Holocene rhyolitic eruption in the Dyngjufjöll volcanic centre and the formation of the Askja caldera, north Iceland. Bull. Volcan., 64, 192–205. Sigvaldason, G.E., Annertz, K., Nilsson, M. (1992). Effect of glacier loading/unloading on volcanism: Postglacial volcanic production rate of the Dyngjufjöll area, central Iceland. Bull. Volcan., 54, 385–392. Singer, B.S., Guillou, H., Jicha, B.R., Laj, C., Kissel, C., Beard, B.L., Johnson, C.M. (2009). 40Ar/39Ar, K-Ar and 230Th-238U dating of the Laschamp excursion: A radioisotopic tie-point for ice core and climate chronologies. Earth Planet. Sci. Lett., 286(1–2), 80–88. Sinton, J., Grönvold, K., Sæmundsson, K. (2005). Postglacial eruptive history of the western volcanic zone, Iceland. Geoch. Geophys. Geosys., 6, Q12009 [Online]. Available at: http://doi.org/10.1029/2005GC001021. Slater, L., Jull, M., McKenzie, D., Grönvold, K. (1998). Deglaciation effects on mantle melting under Iceland: Results from the northern volcanic zone. Earth Planet. Sci. Lett., 164, 151–154. Smellie, J.L. (2000). Sub-glacial eruptions. In Encyclopedia of Volcanoes, Sigurdsson, H. (ed.). Academic Press, San Diego.

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Smellie, J.L., Mcintosh, W.C., Esser, R., Fretwell, P. (2006). The Cape Purvis volcano, Dundee Island (northern Antarctic Peninsula): Late Pleistocene age, eruptive processes and implications for a glacial palaeoenvironment. Antarctic Sci., 18(3), 399–408. Smellie, J.L., Rocchi, S., Johnson, J.S., Di Vincenzo, G., Schaefer, J.M. (2018). A tuff cone erupted under frozen-bed ice (northern Victoria Land, Antarctica): Linking glaciovolcanic and cosmogenic nuclide data for ice sheet reconstructions. Bull. Volcan., 80(12) [Online]. Available at: http://doi.org/10.1007/s00445-017-1185-x. Stevenson, J.A. (2004). Volcano–ice interaction at Öraefajökull and Kerlingarfjöll, Iceland. PhD thesis, The Open University, University of Lancaster, Lancaster. Tomlinson, E.L., Thordarson, T., Müller, W., Thirlwall, M., Menzies, M.A. (2010). Microanalysis of tephra by LA-ICP-MS – Strategies, advantages and limitations assessed using the Thorsmörk ignimbrite (Southern Iceland). Chem. Geol., 279, 73–89. Van Vliet-Lanoë, B., Bourgeois, O., Dauteuil, O., Embry, J.C., Guillou, H., Schneider, J.L. (2005). Deglaciation and volcano-seismic activity in Northern Iceland: Holocene and early Eemian (The Syðra Formation). Geodin. Acta, 18, 81–100. Van Vliet-Lanoë, B., Bergerat, F., Allemand, P., Innocent, C., Guillou, H., Cavailhes, T., Liorzou, C., Grandjean, P., Passot, S. (2020a). Tectonism and volcanism enhanced by deglaciation events in southern Iceland. Quat. Res., 94, 94–120 [Online]. Available at: http://doi.org/10.1017/qua.2019.68. Van Vliet-Lanoë, B., Knudsen, O., Guðmundsson, A., Guillou, H., Chazot, G., Langlade, J., Liorzou, C., Nonnotte, P. (2020b). Volcanoes and climate: The triggering of Preboreal jökulhlaups in Iceland. Int. J. Earth Sci., 109, 847–876 [Online]. Available at: http://doi.org/10.1007/s00531-020-01833-9. Werner, R. and Schminke, H.-U. (1999). Englacial vs lacustrine origin of volcanic table mountains: Evidence from Iceland. Bull. Volcan., 60, 335–354. Werner, R., Schmincke, H.-U., Sigvaldason, G. (1996). A new model for the evolution of table mountains: Volcanological and petrological evidences from Herðubreið and Herðubreiðartögl volcanoes (Iceland). Geol. Rundsch., 85, 390–397.

3

Cenozoic Evolution of Iceland and the Cryosphere Brigitte VAN VLIET-LANOË and Hervé GUILLOU

3.1. Ice ages and the opening of the Atlantic At the beginning of the Tertiary, there did not seem to be any evidence of the existence of ice masses in polar zones, except locally in East Antarctica on the reliefs. Exchanges between the Arctic and the Tethys ocean took place via the Eurasian platform through the Barents Sea (Vorren et al. 1991) and the Turgaï Pass, east of the Urals: the Arctic basin had been lacustrine since the Jurassic and bordered by forests until the dawn of the Eocene (Figure 3.1). The Bering Strait was still non-existent. Continental platforms still linked Europe to Greenland, in the Scoresbysund sector (Beard 2008; Ellis and Stocker 2014) due to the existence of a very large peneplain that was formed by steps since the Variscan orogeny. Elements of it are found perched in eastern Greenland (Japsen et al. 2014; Døssing et al. 2016). The installation of the Antarctic glaciation formally began at the end of the Eocene, ca 40 million years ago, as mountain glaciers on the Trans-Antarctic Range and the East Antarctic Mountains, probably already frozen since the Cretaceous (Figure 3.2). Nevertheless, it occurred mainly at the end of the Oligocene (30 My), as a

For color version of the figures in this chapter see www.iste.co.uk/vanvliet/iceland2.zip Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021. Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

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result of the increasing oceanic isolation of Antarctica after the separation of India and Australia.

Figure 3.1. Arctic paleogeography in the final Lower Miocene (modified from Jakobsson et al. 2007). Iceland had not yet emerged. Is: location of submerged Iceland, in white; GIFR: Greenland-Faroe Ridge; ÆR: Ægir Paleoridge; GR: Greenland Ridge; YP: Yermak Plateau

The glaciation of the Northern Hemisphere seems to have occurred later, due to (1) the presence of a giant lake in polar position (Figure 3.1); (2) that of temporary sea and lake ices as early as the Maastrichtian (65 My) (MacLeod et al. 2011), then seasonally since 47 My in the Arctic basin; its size increased during the Upper Eocene from 40 My, allowing the formation of deep salty water in the Northern Atlantic (Davies et al. 2001); and finally (3) the formation of coastal reliefs that was still little marked before the Eocene. The Fram Strait (Figures 3.1 and 3.3(a)) in the north Atlantic began to open (Figure 3.3(b)), but was still insufficiently deep (– 2,500 m), leading to a first significant superficial inflow of seawater traced by the arrival of foraminifera in the Arctic basin from the Middle Eocene at about 40 My (Moran et al. 2006). Exchanges occurred mainly via the Barents Platform covered by shallow sea (< 200 m). Zonal oceanic circulation disappeared with the closure of the Neo-Tethys Sea at about 35 My (Von der Heydt and Dijkstra 2006) due to the collision of the Arabian plate with Eurasia. The Greenland Sea opened a little more at about 25 My, in association with a reorganization of rifting in the North Atlantic at about 27 My (Chapter 3 of Volume 1). The Arctic Lake became distinctly brackish.

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Figure 3.2. Thermal and eustatic events leading to the onset of glaciation (modified and completed from Zachos et al. 2008). Basaltic effusive phases (black bars) in sedimentary basins are extracted from (Ernst and Buchan 2002)

The first evidence of glaciation in the Northern Hemisphere was observed in the form of IRD (ice-rafted debris released by icebergs) in marine sediment cores in the

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Northwest Atlantic. These results suggest the onset of glaciation in eastern Greenland either about 45 My (Moran et al. 2006; Tripati et al. 2008) or about 38–30 My, in synchrony with the onset of Antarctic glaciation (Eldrett et al. 2007, 2009) (Figure 3.2). The observation of rolled gravels at 45 My suggests a drift by the sea ice of fluvio-glacial inputs from the nascent landforms as proposed by Stickley et al. (2012) rather than a direct glacial input. Mountain glaciers may have existed since 35 My on the Greenland Coastal Mountains in the process of uplifting (Japsen et al. 2014) and on the Alaskan Coastal Cordillera, as shown by the sedimentation of the Seward Peninsula in Alaska. The appearance of true calving glaciers is dated between 38 and 30 My (Bernard et al. 2016) and demonstrated by the occurrence of IRD found more than 300 kilometers east of the Greenland coast (Figure 3.3(a)). Their petrography is similar to that of the walls of the fjords of East Greenland, including the Scoresbysund, between 68° and 76°N. This phase corresponds to a phase of rapid exhumation of the East Greenland margin and the Scandinavian margin (Japsen et al. 2014). The major cooling at ca 38 My is associated with a decrease of about 4–5°C in the temperature of deep ocean waters and about 10°C in the average temperature of high latitude land masses; sea ice reappeared more frequently in the North Atlantic (Stickley et al. 2012). These features also imply a stepwise modification of the ocean floor due to the northward extension of the Atlantic rifting, as suggested by indications of volcanic activity during the Miocene on the Sverdrup Bank (Yermak Plateau, located immediately north of the Fram Strait; Figure 3.3) (Geisseler et al. 2011) and the probable emergence of Iceland ca 25 My ago (Hjartasson et al. 2017) as a volcanic plateau overlying an ancient basalt substratum (> 54 My) (see Volume 1, Chapter 3). Its oceanic isolation occurred at about 15 My, following the subsidence of the Greenland–Iceland Ridge. This subsidence was associated with the uplift of the Scandinavian margins (Japsen et al. 2014), which became high enough to allow the formation of glaciers or ice caps (Figure 3.3(b)). The Arctic Ocean was salty and well ventilated at about 17.5 My, attesting to a true opening of the Fram Strait (Jakobsson et al. 2007). Nevertheless, its salt content remained significantly lower than that of other oceans. Sedimentation on the Greenland–Iceland–Faroe Ridge and on the northern margin of the British Isles changed in style from the Late Oligocene and especially the Early Miocene, with the overflow of cold bottom currents to the east and on the threshold of the Faroe Islands (Figure 3.3(a)). The thermal subsidence of the continental shelves (Rockall Plateau) and the troughs south of the Greenland-Faroe Ridge (from 500 m for the banks to 2,000 m for the troughs) allowed the recording of changes in deep ocean currents since

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35 My, especially during the Lower Miocene (15 My) and the Early Upper Miocene (9 My, Tortonian), resulting in an ocean circulation close to the present one (Davies et al. 2001; Hutchinson et al. 2019). In addition, due to water storage in the ice caps and oceanic thermal retraction, there was also a stepwise lowering of the interglacial sea level since about 35 million years ago: this is the complex regression of the Oligocene, from +150 m to 0 m for warm periods to today.

Figure 3.3(a). Location of sites. In red: current North Atlantic Deep Water (NADW) crossings. Image reproduced from GEBCO, 2014. Processing M. Jaud

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Figure 3.3(b). Tectonic and hydrological evolution of the North Atlantic since the Eocene. GIR: Greenland–Iceland Ridge; IFR: Iceland-Faroe Ridge; NADW: North Atlantic Deep Water. Margin uplift phases from (Japsen et al. 2014). Rifting velocity in Southeast Greenland from (Ellis and Stocker 2014) (column width: 2.5 cm/yr)

3.1.1. The Middle and Final Miocene cooling Traces of glaciation appeared more regularly from the beginning of the Miocene, on the edge of the continental shelf, following the important Save orogeny, interrupted by brief warming, and then become more pronounced at about 14 My, then especially at the end of the Miocene. At this time, IRD were observed in North Atlantic cores, showing calving glaciers and iceberg drift (Bleil 1989) (Figure 3.4). On the Vøring Plateau, southwest of the Lofoten Islands, the first traces of IRD appeared at about 12.6 My, indicating that Scandinavian glaciers reached the sea during the Middle Miocene.

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The widening and deepening of the Fram Strait were marked as early as 9.8 My and especially after 6–5 My, promoting the accentuation of thermohaline circulation and the cooling of the North Atlantic system during the Neogene and even the Middle Quaternary (Hodell et al. 2008). The opening by subsidence of the Bering Strait became effective at about 6 My. The functioning of the Møhn volcanic ridge, north of Jan Mayen Island (Géli 1991), slowed down by 50% from 9 My to 6–5 My. The subsidence of the shallows of the Greenland-Faroe Ridge increased during the Middle Pliocene, then accelerated again at ca 3–2 My, with an increase in tectonovolcanic activity. The latter corresponded in Iceland to the last rift jump. An extensive ice cap over Greenland has existed since 9 My (Larsen et al. 1994; Solheim et al. 1997) and the winter ice sea went down to the Vilaine estuary in Brittany (6.7 Ma) (Van Vliet-Lanoë et al. 2010), almost synchronously with a very significant glaciation in Antarctica (6.7 My) (Herbert et al. 2016).

Figure 3.4. Chronology of the onset of ice-rafted debris in the North Atlantic. From 600,000 years onwards, their occurrence was even more significant (off the scale in the figure; from (Bleil 1989)). MPT: Middle Pleistocene (Climatic) Transition

3.1.2. The acceleration of the Middle Pliocene 3.1.2.1. The Middle Pliocene During the Late Middle Pliocene (3.6–3.3 My), when the climate was relatively “mild” and the last passages from the Pacific to the Atlantic were still open in Panama, two small ice caps persisted over Southern and Central-eastern Greenland

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(De Scheppers et al. 2013). With the Bering Strait open, evidence of glaciation was becoming more frequent on the Barents Shelf and inland. A very important and isolated glaciation developed at about 3.2 My (isotopic stage M2; 3.26–3.02 My); its overload probably triggered in Iceland the last rift jump at 3 My. Sea ice appeared regularly in the Icelandic Sea and probably promoted the extension of the Greenland ice sheet (Clotten et al. 2018). The ice caps became widespread during glacials at the end of the Pliocene, that is after 3.2 My (Figure 3.5). A gradual decrease in average sea level (first order) accompanied this new step towards glaciation. Further intensification of North Atlantic coastal glaciation was observed at ca 2.7 My (Jansen et al. 2000; Kleiven et al. 2002; Haug et al. 2005; Bartoli et al. 2006). From 2.8–2.4 My, small ice sheets appeared regularly in the northern hemisphere. The onset of the great Quaternary glaciations is dated to 2.75 My on the Vøring Plateau, and to 2.9 My on the Icelandic Plateau, northwest of Iceland. This shift reflects a difference in chronology and growth of the ice caps between the Greenlandic margin and Scandinavia, in relation to thermohaline circulation. This intensification was also noticeable during sedimentation on the Yermak Plateau (Vogt et al. 1994; Geissler et al. 2011). It was marked by a significant contribution of IRD between 2.7 and 2.4 My and also by an increasing duration of sea ice from 2.7 to 2.1 My. The cyclicity of the glaciations was then rapid, close to 41 ky, corresponding to the frequency of oscillation of the Earth’s rotation axis (obliquity), and thus to a North Pole triggering of the glaciations. 3.1.2.2. The Pleistocene Cooling really increased ca 2.3 My ago, possibly related to the explosion of a supernova (Breitschwerdt et al. 2016). This marked the beginning of the Pleistocene. The record of IRD transported by icebergs attests to a persistent plateau glaciation (fona or small ice cap) over Scandinavia and fresh interglacials, between 2.5 and 1 My, associated with a dramatic increase in glacial erosion. The maximum extent of the North American and Scandinavian ice sheets (Figure 3.5) as well as that of the Alps is recorded at 2.2 My. It was related to a still high sea level and a functional Gulf Stream, as shown by the often still temperate character of Scandinavia, whose maximum glaciation occurred much later in the south, due to the presence of the mild North Atlantic drift, that is the continuation to the North of the Gulf Stream. Between 2.1 and 1.2 My, however, there were no large ice caps. Glaciations remained limited over Scandinavia and conditions in the Norwegian Sea demonstrated a moderately cold context with periods of limited influence of Atlantic waters and a significant extent of winter sea ice, while others periods evidenced the penetration of Atlantic waters into the eastern Arctic Ocean. Palynological analysis of the glacial deposits of this period shows that it was no colder than the Holocene, our present interglacial.

Figure 3.5. Evolution of peri-Atlantic freezing and thermohaline circulation in relation to the activity of the North Atlantic ridges in gray (Fram; Jan Mayen, Iceland). δ18O: (Lisiencki and Raymo 2005). NADW: North Atlantic Deep Water (current). NCW: deep cold water inflow. ASW: Arctic surface water inflow into the North Atlantic Basin. MPT: middle Pleistocene Transition (Van Vliet-Lanoë, unpublished)

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After 1 My, the contrasts between glacial and interglacials became more pronounced, due to a greater extension of the Scandinavian ice cap and an acceleration of the North Atlantic drift during interglacials. However, as Ruddiman and Mac Intyre pointed out as early as 1971, ice caps form when sea levels are still high. Ice streams (section 3.21) appeared in Greenland and the Barents Ice Sheets (Hogan et al. 2016). Along the west coast of Norway, near Bergen and Trondheim, the Fedje glaciation (Norway) was the first period of glacial erosion eroding Tertiary sediments from the continental margin; it has been correlated with the Menapian glacial (1.1 My, marine isotope stage or MIS 34). 3.1.3. The Middle Pleistocene Transition Thereafter, from 1.1 to 0.6 My, a period also known as the Middle Pleistocene Transition (MPT, Figure 3.5), glaciations were less extensive and rarely reached the edge of the continental shelf, such as on the Yermak Plateau in the Barents Sea at about 950 ky. Their extent was about two-thirds that of the largest Pleistocene ice sheets, and interglacials were also less pronounced than those of the Late Pleistocene, particularly MIS 31, which was extremely warm, even in the Arctic (HernándezAlmeida et al. 2013). This climatic episode corresponded to an important activity of the Icelandic hotspot and the North Atlantic rift. The sea ice was spreading in the Greenland Sea and icebergs from Greenland drifted towards mid-latitudes with the widening of the Fram Strait. The ice caps, which were initially located at high latitudes during the Pliocene, were refocused towards 70–50° N. This episode ended with the major glaciation of MIS 16 (680 ky). The discovery of distinct glacial striations and scourings on the surface of the continental shelf bordering the Arctic Ocean has opened a debate on a complete freeze up of this ocean in the manner of the margins of the Antarctic ice sheet (formation of ice shelves). This idea had already been put forward by Hughes et al. (1977). An ice margin can float from 500 m thick up to 2,000 m thick. These observed striations appear at a depth of 1,000 m: – on the Lomonosov Ridge, in the center of the Arctic basin; – on the Chukchi Peninsula bordering the Siberian Platform and the Bering Strait; – at 400 m depth on the south of the Yermak Plateau, north of the Fram Strait; – on the Peary Shelf in northern Greenland (Figure 3.6). The deepest ones are probably related to a Neogene subsidence, especially in the center of the basin and on either side of the Gakkel Ridge.

Figure 3.6. Two hypotheses of glaciation during the Weichselian Pleniglacial. (A) Complete freezing hypothesis (modified from (Hughes et al. 1977)). AIS: Arctic ice cap; BRIS: Barents ice cap; BIS: British Isles ice cap; SIS: Scandinavian ice cap; GIS: Greenland ice cap; LIS: Laurentide ice cap; (B) A more plausible version of partial glaciation during the same period with ice rivers and ice shelves (Bjarnadóttir et al. 2014) LR: Lomonosov Ridge

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The development of floating ice margins or ice shelves in the Arctic Basin is currently being studied (Bjarnadóttir et al. 2014; Jakobsson et al. 2014). These submarine plateaus and ridges are thermally subsident and these features could be attributed either to one of the major glaciations of the Mio-Pliocene (6.7 My, 3.2 My) or Quaternary (MIS 16: 680 ky, MIS 12: 450 ky), or even more locally to MIS 6 or to the last glaciation, the Weichselian. The complete glacierization of an ocean basin surrounded by continental masses is very difficult to justify from the point of view of precipitation, given the absence of evaporation due to the continuous nature of the sea ice. Taking into account the deformation location map (Figure 3.6), it seems more plausible that some local platforms produced these figures, for example on the Lomonossov fossil Ridge. Real icebergs 1,000 m high (glacial) included in the sea ice can drift across the basin, under the impact of strong winds and the Arctic marine gyre, and cross the top of the ridge. 3.1.4. The initiation of thermohaline circulation The Icelandic Sea, located on the deep Icelandic Shelf, is a critical site for recording climatic variations affecting abrupt changes in the paths of shallow ocean currents (Figure 3.6, section 3.4.2) and the glacial evolution of Greenland, and thus the evolution of oceanic circulation and climate since the Neogene, via thermohaline circulation (THC). The position of Iceland, only surrounded by ocean since about 15 My (section 3.4.3.1), also makes it a key witness to the changes controlling the geodynamic evolution of the North Atlantic via isotopically dated marine and continental records (Chapter 2). The presence of an ice sheet/shelf or sea ice in the North Atlantic results in an active and deep ocean circulation, increasing the interzonal thermal gradient, the equivalent of our present-day thermohaline circulation. The extension of the Arctic Sea ice is primordial and synchronous with the Late Eocene glaciation of Antarctica and probably of Greenland (formation of the coastal reliefs). With the northward drift of India and Australia, the role of thermal drive played by the Southern Ocean (intertropical ocean) gradually increased to modulate the glacial-interglacial system that had been developing since 40 My. It interacted notably by forcing thermal events in the northern hemisphere such as those of Bond (Holocene and interglacials) or warm Dansgaard–Oeschger (DO) during the Ice Age. The first evidence of deep water circulation in the trough located east of the Faroe Islands trough appeared about 35 My ago (Cramer et al. 2009; Dixon et al. 1994), as we have seen previously (section 3.2.2.3 of Volume 1), as a result of

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thermal subsidence along the Iceland-Faroe Ridge in the extent of the Cretaceous rift and Rockall Basin, characterized by a thinned oceanic crust. The subsidence of the Iceland-Faroe Ridge (IFR) was early and powerful (35 My) and the Oligocene submarine abrasion platform is today much deeper than in the west of the island. The thermal subsidence of the Greenland–Iceland Ridge (GIR) was significantly lower during the first phase (35 My) and reached a maximum of only 300 m, due to the migration of the hotspot as the ocean opened. Analysis of the concomitant uplift of Greenland and Scandinavia, dated via fission tracks (Japsen et al. 2014), show that these marine subsidence events occurred synchronously with stages of uplift and cooling of continental margins: during the Upper Eocene/Early Oligocene (between 40 and 35 My); during the Miocene, about 15 My ago; and finally during the Pliocene (at ca 5 My). This implies a decoupling between uplifting continents and passive margins, but also a coupling between cooling of the ocean mass and GIR subsidence. The opening of the Fram Strait during the Late Eocene (38–35 My) and the activation of the Mid-Atlantic Ridge following the tectonic reorganization of the North Atlantic connected the Arctic Lake with the world ocean, and triggered the thermohaline circulation as the zonal oceanic circulation stopped within the Tethys. The Greenland Sea opened a little more, in association with a transtensive reorganization of the rifting in the North Atlantic at 27 My (see section 3.2.2.3 of Volume 1). Thermohaline circulation was already functional in the Norwegian Sea 23 My ago, and most of the circulation was still restricted to the eastern Atlantic Ocean via the Rockall Trough until 20 My: the already subsiding ridge between Iceland and the Faroe Islands became flooded (Uenzelmann-Neben and Gruetzner 2018), initiating the Iceland-Scotland overflow water (ISOW). In the west, shallow waters prevailed with sometimes large islands or paleolakes between proto-Iceland and Greenland: these were the continental bridges of botanists. Thereafter, the ridge between Greenland and Iceland collapsed ca 15 My ago, leading to the complete insularization of the island. Despite the absence of land in the Arctic, the effective start of thermohaline circulation at the beginning of the Neogene allowed for the thermal homogenization of the system via the inertia of the oceanic mass, while the thermally isolated South Pole remained cold. Current thermohaline circulation was gradually taking place, especially after the opening of the Bering Strait and the progressive closure of the Balboa Strait (Panama isthmus), also about 15 My ago (Montes et al. 2012), blocking the last inputs of warm Tethysian waters that disappeared between 16 and 15 My (beginning of the Serravallian). The Isthmus of Panama discreetly emerged then, 6 My ago. The significant influx of Pacific waters into the Arctic Basin was related to the deepening by subsidence of the Bering Strait, allowing the arrival of Pacific mollusk

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larvae until about 4.5 My in the Tjörnes Formation in Iceland (De Scheppers et al. 2015) (section 3.2). In conclusion, the morphology of the North Atlantic basin and its connections to the Arctic and Labrador basins are controlled by plate tectonics. For this reason, the current low contribution of the deep water from the Arctic Basin to the global thermohaline circulation limits interlatitudinal heat exchange. If the future development of the Mid-Atlantic Ridge favors a deepening and widening of the Fram Strait, the thermal regime of the world ocean could be severely disorganized. 3.2. Iceland’s Quaternary glaciations 3.2.1. Conditions for the development and functioning of ice caps The development of glaciations in Iceland is closely related to (1) the extent of sea ice, (2) precipitation on land under the control of heat input from the Irminger Current (Figure 3.7; see section 3.4), and (3) the landforms inherited from the volcanic activity of the hotspot. This relief promotes the formation of the ice cap, under the direct influence of the cooling brought by the East Icelandic current or drift, itself derived from the East Greenland current. It is for these reasons that the formation of the ice sheet occurs in the south-east of Iceland, where the precipitation and the topography are today the most important. The north of the island is now relatively dry (section 3.4) and generally arid during cold periods, which is a limiting factor for local glacial extension (Figure 3.7). Ice masses are transiting northward from the current glacial centers. During the thermal optimum of interglacials, glaciers generally disappear, but are reconstituted during cooling, especially close to the end of the interglacial periods. Because of its position in the central North Atlantic Ocean, Iceland acts as an intergrade between Greenland and East America, cooled by the East Greenland Drift, and Scandinavia, warmed and watered by the North Atlantic drift (see Figure 3.32). During deglaciation or early interglacials, the mild Irminger Current was dominant (Jennings et al. 2000), especially on the west coast (Figure 3.7). On the other hand, at the end of the interglacial or during glaciation, cold currents dominated with often a late warming. The presence of sea ice limits the vaporization of the ocean and thus the accumulation of snow on land and the feeding of glaciers. The Icelandic ice cap has

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never formed a permanent mass like those of Greenland or Antarctica. During cooling episodes, the ice cap is generally cold in altitude (on permafrost) and thickens like those of the Western Russian Arctic. It has a rather domed morphology during glacials.

Figure 3.7. Relationships between high present precipitation (gray), position of ice caps and ocean currents. When most of the Irminger Current veers westward, during cooling, precipitation is very low in the north and limit the formation of caps (white bars). The extension of the sea ice on the northern part of the island plays in the same direction

During climate warming, ice overload due to high precipitation causes a lowering of the melting point at the base of the ice cap and a basal lubrication of the glacier feeding area, while at lower altitudes the plasticity of the ice increases with temperature. Also at low and medium altitudes, the melting of snow in summer and the overall increase in precipitation also leads to lubrication and sustained hydrostatic pressure at the base of the glacier, with water often being evacuated through sub-glacial drains (eskers, valley-tunnels, fractured bedrock). Outburst dynamics or surge (Figure 3.8(A)) thus appears in coupling with an acceleration of ice cap outlets or ice streams, which are very fastflowing glaciers

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(100 m/year to 100 m/day). Currently, the majority of Icelandic glaciers are of the ice stream or surging types (Bjornsson 2017; Björnsson et al. 2003). The surface of a surging glacier may subside by a few tens of meters as a result of rapid drainage (a few days) under hydrostatic load of a subglacial or intraglacial aquifer (Figure 3.8(B)). The thermally increased plasticity of the ice will induce a gravitational spreading of the ice sheet (Figure 3.8(D)), which takes on a flattened profile and is drained by ice streams. If the outburst reaches the coast, a glacial tongue detaches from the substratum and floats. It will very quickly invade the fjords, then advance in line with the shelf, coalesce and spread out in an ice shelf (Figure 3.8(C)) or a floating ice shelf like the past Disko ice shelf in West Greenland.

Figure 3.8. Ice streams and ice shelves

COMMENT ON FIGURE 3.8. – (A) Surging Brögger Glacier (downstream part; northwestern Svalbard; Norsk Polar Instytutt Photo Library). (B) Surging glacier with collapse of a single outlet (southeast Svalbard) (© Brigitte Van Vliet-Lanoë). (C) Rectilinear surge tongue of the Drygalski glacier through the Ross Sea ice shelf (Antarctica, photo Earth Observatory, NASA). (D) Freeman glacier (Edgeøya,

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Svalbard) in gravity spreading, after its 1956 surge in a zone warmed in summer by a branch of the North Atlantic drift (photo Norsk Polar Instytutt). The floating glacial tongues contribute to the rise of the sea level and can spread at the sea surface in ice shelves, as in Antarctica on that of Amery in the east, confluence and flow, as on that of Ross Sea. Nevertheless, small ice shelves may form without an ice sheet upstream, as on Northern Ellesmere Island (Canada) (Vincent et al. 2001) or in Antarctica, on the western peninsula, where the Larsen B platform was dismantled in 2002. This dismantling phenomenon is generally initiated when the temperature rises again. This is observed during or after Dansgaard–Oeschger (glacial) or Bond (Holocene) events in the North Atlantic under the impact of periods of high solar activity (Rasmussen et al. 2016; Boer et al. 2018). In the context of a stadial or cold episode in a glaciation, these floating ice shelves may last a few thousand years. Abundant calving of floating ice strips, or ice shelves, will lead to the arrival of icerafted debris (IRD) in ocean sediments. Nevertheless, a surge accompanied by the rapid calving retreat of an ice front does not necessarily prejudge global warming. It may be only linked to the functioning of its watershed, with periods of inertia, corresponding for example to the charging of the glacier internal aquifer. The Vatnajökull glaciers in Iceland were all surging during the warming following the Little Ice Age. These surges were controlled both by the rise in temperature and by the increase in precipitation linked to the return of the predominantly positive character of the North Atlantic climatic oscillations NAO and AMO (North Atlantic Oscillation and Atlantic Meridional Oscillation) or the South Atlantic Circulation, transmitted by the Irminger Current especially during Dansgard-Oeschger warming events (van Kreveld et al. 2000). Finally, the untimely advance of a surging glacier on land can lead to the formation of large dam lakes, which may cause a flash flood or jökulhlaup (section 1.5). 3.2.2. Glacio-isostasy 3.2.2.1. The loading of the earth’s crust The loading induced by ice sheets leads to a progressive sinking of the earth’s crust, in other words a glacio-isostatic subsidence. This flexure is slow in the context of thick continental crust (> 35 km) and therefore rigid (old Archean shields). This is the case of the Canadian and Scandinavian shields. It is much faster in the case of a young, relatively thin and warm crust such as that of Southern Iceland. These

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isostatic readjustments are compensated by mass creeps within the lower, more viscous lithosphere. The amplitude of the deformation can exceed 900 m. Currently, due to the glacial load, the center of Greenland (ice thickness > 3,000 m) is depressed and its bedrock has sunk below the current sea level. Mass rebalancing during deglaciation can lead to rapid uplift, exceeding 15 cm/year, in the case of thin crust (Spitsbergen, Iceland). For Southern and Western Iceland, values of 8.8–10.5 cm/year have been estimated (Biessy et al. 2008; Ingólfsson et al. 2010; Le Breton et al. 2010). The end of the uplifting is reached in 1,000 years in this case, whereas it can take more than 10,000 years in the case of thick crust and is still not complete for Scandinavia and the northeastern part of the Canadian Shield. In the case of large ice caps, the readjustment lasts several thousand years due to the volume of displaced mantle: the load compensation depth is of the order of one kilometer, which in Iceland represents the depth of the shallowest magmatic reservoirs. In the case of regional ice caps, the load is only compensated by a superficial and rapid flow of the asthenosphere, the subsidence of the crust being very local and limited. A thick, old and rigid crust, such as a piece of continental crust, will react more slowly than a thinner crust, such as an oceanic crust. This explains pro parte the partitioning of the Early Holocene isostatic rebound along the Icelandic coast (Le Breton et al. 2010) and of course conditions the thickness of the ice cap and the history of its deglaciation. This temporary subsidence of the crust, as well as the presence of morainic ridges, confines drainage to the vicinity of volcano-glacial structures, leading to the formation of very large lakes, juxtaposed or sub-glacial (glacier aquifers) during the deglaciation period. These lakes are often included in external morainic arcs, which, when they burst, create large flash floods (jökulhlaups) (section 1.4.3). An uplift rate of up to 3 to 4 cm/year can be observed in the central part of the Vatnajökull ice cap (Sturkell et al. 2003), and is associated to its current deglaciation. This very rapid uplift favors an incision of the proglacial river systems with a downstream evacuation of sediments, also promoted by the increased evacuation of meltwater and recurrent jökulhlaups. 3.2.2.2. Glacio-eustatic implications During periods of global warming, the evolution of relative sea level will be the combined result of the sea level rise related to ice melting and the glacio-isostatic rebound of the basement (Figure 3.9).

Figure 3.9. Proglacial reaction flexural bulge: the deformation propagates in front of the ice cap (h2), but with less intensity than at the core of the cap (h1). The reaction bulge will migrate marginally and centrifugally as the cap extends. The same dynamic occurs in the opposite direction during deglaciation. Graph B shows the changes in the modalities of postglacial readjustment during deglaciation of an ice sheet (from Andrews 1970)

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Near the ice caps and on thin crust, the glacio-isostatic adjustment dominates, as in Iceland. On the other hand, at a greater distance from the ice centers, it is the load of the global Holocene eustatic transgression linked to the melting of the Antarctic and Greenland ice caps, coupled with the collapse of the reactionary marginal forebulge of the upper crust, that dominates, as presently in the Southern North Sea and Northern Germany. Today, the global sea level is still rising gradually by 0.3 mm/year due to the uplift of polar zones that have not yet reached glacio-isostatic equilibrium (glacio-isostatic accommodation of models or GIA). In the longer term, after rebalancing and complete melting of all the world’s ice caps, this global uplift could account for the 150 m of eustatic lowering recorded over the last 20 million years. The Holocene transgression is limited in Iceland to a eustatic rise of about 75 m due to the complex history of the subsidence of the basement, kept depressed by Holocene variations in the ice cap. Thus, we observe first of all a submersion at the beginning of the glaciation with sedimentary facies, first lacustrine then marine (Late Glacial), followed by re-emergence during the Younger Dryas (glacial advance). It produced a renewed subsidence lasting until the Pre-boreal about 9,500 years cal BP ago, followed by re-emergence of the coast line met by global transgression at ca 6,000 years cal BP (Ingófsson et al. 2010). It submerged the lava flows emitted during the peak of volcanic emissions concomitant with the deglaciation (ca 12–8 ky; Figure 3.10).

Figure 3.10. Apparent sea-level rise curve for the west coast of Iceland (completed from Ingólfsson et al. 2010)

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COMMENT ON FIGURE 3.10. – It is important to note that the maximum apparent regression corresponds to the peak of effusive volcanism related to deglaciation. Term Ia: Termination Ia (Bølling deglaciation); Term Ib: Termination Ib (Holocene transgression). Relative uplift is linked to temporary glacial loading (depression). 3.2.2.3. Sedimentary implications in sequential stratigraphy During glaciation, the facies recorded will most often be bottom till or basal till, highly compacted, corresponding to glacial shear, active glacio-tectonics and subsidence. They may be preceded on land by thin fluvio-glacial formations or, in the littoral zone, by iceberg scours or released deposits and stepped littoral formations, contemporary with the eustatic fall. The subsidence by glacio-isostatic loading, when it is relatively fast and follows the eustatic lowering, will also limit the growth of the ice caps by promoting synchronous calving of the marine outlet glaciers from the ice cap. During deglaciation, the scouring phenomena by eskers or tunnel valleys are important, due to the formation of an important aquifer, internal to the ice cap; the nature of the deposits is more varied (kame, ablation till, filling of meltwater basins, varves), often accompanied by passive or active glacio-tectonics in relation to temporary re-advances. These facies are clearly visible in the Skjálfandafljót valley downstream of Laugar (Northern Iceland) and for glacio-marine facies northeast of Vopnafjörður and in Breiðavík Bay (Tjörnes Peninsula). 3.2.3. Icelandic data 3.2.3.1. Specifics Curiously, the Pleistocene stratigraphy of Iceland is not very detailed, even if at least 20 glaciations were recorded for the entire Quaternary based on K-Ar datings (section 2.2) for lavas older than 1 My and paleomagnetic data for the most recent. While the Early Pleistocene is well documented on land, the stratigraphic data for most of the Quaternary are limited in time (very few available ages) and space, except from the Younger Dryas (13-12 cal ky). The last deglaciation has been dated essentially by radiocarbon on shells, in marine cores drilled on the platform of the island, or on land in the outlet zones of the great fjords, which sometimes poses problems of volcanic contamination or carbon reservoir age. Many stratigraphic controversies are related to the limited possibilities of dating, especially for the Middle and Upper Pleistocene, because of the reduced accessibility of outcrops on land. The primary sources of chronological data in Iceland are (1) magnetic polarities (normal or reverse) of lavas, (2) K-Ar and more

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recently 40Ar/39Ar datings (section 2.2), and (3) tephrostratigraphic data (section 1.4.4.1) for the Upper Pleistocene and Holocene, coupled with the geochemistry of major elements, and, since the 1990s, trace elements (section 1.2).

Figure 3.11. (A) Indications of moderate hydrothermal activity in Holocene sediments of Fossvogur (Reykjavík) (© Brigitte Van Vliet-Lanoë). (B) Heavily reddened fracture in the basalt at Ytri Vík (photo © ISOR)

Radiocarbon dating (14C) poses several problems in Iceland. As seen in box 2.1, the range of use of this method corresponds to absolute ages from a few hundred

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years to, at most 50–60 ky. Some deposits dated in this time range are likely to belong to the previous interglacial. On the other hand, integration of atmospheric carbon by marine waters takes some time and the global ocean circulation is a long cycle (about 1.5 ky long) that includes deep currents whose upwelling brings to the surface waters several hundred years old. These phenomena imply that the carbon in ocean waters is on average 400 years older than that in the atmosphere (global reservoir effect). However, when the ice caps melt, the CO2 included in the vesicles contained in the ice will also affect the organisms calcified during this period. The vast majority of 14C ages at high latitudes are obtained on marine shells, with a non-negligible reservoir effect (R, with a correction expressed as δR), generally corresponding to a value of δR of 400 years, as suggested by Walbroeck et al. (2008). But a more precise analysis of deglaciation periods shows very unstable δR values due to melting of ice caps, with δR values often exceeding 1,000 years during the melting period (a few decades) of ice of various ages. In the Icelandic context, fumarolic emissions can locally enrich the atmosphere and waters with ancient carbon from magma or ancient organic reservoirs (hydrocarbons, such as in the Faroe Islands or Greenland). The presence of calcite recrystallization and shrinkage patterns in the sediments is evidence of hydrothermal activity and therefore invalidates their dating (Figure 3.11). On the other hand, 14C datings older than 43 ky most often attest to shells that are much older, reworked from interglacial deposits, or even earlier glaciomarine deposits. Atmospheric nuclear tests have considerably enriched the atmosphere in 14 C. Since the cessation of these tests, the atmospheric decay of 14C has been very regular, which allows more precise dating of elements younger than 1960. 3.2.3.2. Volcanic tracers of ice volume Altitudinal data from subglacial volcanoes (tuyas) are generally used as a tracer of the thickness of the ice cap and its maximum extension during the Last Glacial Maximum (LGM) (Walker 1965), as these volcanoes are considered Weichselian by many authors, although their K-Ar ages range from 750 ky to the LGM. The altitude of many tuyas is of the order of 1,000 m, taking into account the topography of the Icelandic plateau and the estimated ice surface (Maclennan et al. 2002), thus attesting the existence of a cap about 1,000 m thick in relation to the central plateau. This thickness would have been reached during the LGM (see below, Figure 3.25), in the region north of Vatnajökull, and decreased north and

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west with 400 to 600 m of ice at the current coast. However, according to Guðmundsson (2000a), the altitude of the tuyas is more related to the piezometric level of the lavas stored in magmatic reservoirs than to the thickness of the ice (section 2.1). Several tuyas with fresh morphology could not be dated. Indeed, the amounts of radiogenic argon in the samples collected from some of these tuyas are so low that they could not be detected via mass spectrometric measurements. This observation suggests very recent ages (Guillou et al. 2010) (section 2.3). These include the Herðubreið in the center of the island, whose effusive summit did not evolve after its emplacement. It attests to an ice thickness of about 1,000 m, far less than the 1,500 m previously mentioned (Ingólfsson et al. 1997). Other tuyas with intact morphology such as the Bæjarfjell in the north (473 m) and the Gæsafjöll also appear very young. The Þórósfell, which dominates the Markarfljót to the north, west of the Mýrdalsjökull, seems a little older because of its more evolved periglacial morphology and would attest to an ice thickness of 800 m. 40

Ar/39Ar and K-Ar dating along with detailed stratigraphic work provide a relatively accurate record of the Middle and Upper Pleistocene, allowing a comparison with that obtained from sea cores recovered around the island and in other areas of the Nordic Seas, but also with that of cores from Greenland continental ice cap. Altitudinal data from dated subglacial volcanoes (tuyas) (Guillou et al. 2010 and unpublished data) could be used as a tracer of the thickness of the ice cap, as proposed by Walker in 1965. The values given in the text above correspond to the difference in altitude between the plateau of the tuya and the original plateau (first number) with an addition of 400 m (second number), corresponding to the height of ice above the tuya (current observations of the Grímsvötn volcano). Apart from the events calibrated by absolute dating and paleomagnetism, the current stratigraphy of the Icelandic Quaternary glaciations attributes most of the glacial formations and erosion to the last glacial episode, the Weichselian. 3.2.4. The Icelandic record 3.2.4.1. The very first glaciations (Mio-Pliocene) Iceland emerged at about 25 My, when its hotspot reached the Mid-Atlantic Ridge (Chapters 2 and 3 of Volume 1), but the oldest aerial lavas currently emerging

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were emplaced ca 16 My ago. They are currently outcropping on the periphery of the island, as is the case of the tabular basalts of Akureyri (15 My old), on the northwestern peninsula or on the northeastern part of Iceland. At that time, the island, which was located close to Greenland, barely emerged and took the full impact of the drift of East Greenland waters. Iceland is therefore a very favorable place to analyze the characteristics of the early glaciations during the Upper Miocene and the Plio-Pleistocene transition. The first glacial indices appeared in Iceland at about 7 My, concomitantly with a significant global cooling at 6.7 My, as intercalated sediments within the lava piles of the east and southeast of the island (Frilðeifsson 1995; Hjartason and Hafstað 1997). The fauna and flora recorded in these sedimentary formations intercalated within the volcanic series attest to a progressive cooling from a subtropical context in the Arctic Eocene, then evolving to warm temperate at about 16 My, especially 12 My ago (end of the Serravallian) (Buchardt and Sı́monarson 2003), and then temperate ca 7 My ago. The more recent phases (Figure 3.12) are mainly recorded by the formation of Tjörnes (4.4–2.5 My) north of Húsavík. Different taxa of the Turgaï flora mainly disappeared between 12 and 10 My (Denk et al. 2005) (section 3.4.3.2). Glaciers developed in the southeast since the Upper Miocene (at ca 9 My), reflecting the intermittent development of an ice cap over the southern (highest) area of present-day Vatnajökull (Geirsdottir and Eiriksson 1994) (Figure 3.12). Temperate flora thus developed in Iceland during the Upper Miocene and boreal flora appeared between 6 and 4 My (section 3.4.3.2). The formation of the tundra dates back to 4 My. In the northeast, tillites were intercalated within the basalts (Jokulsá á Brú Valley) since 5 Ma. A drastic change occurred with the glaciation during the isotopic stage M2 (3.26–3.02 My); its glacio-isostatic load probably promoted the onset of the last rift jump 3 My ago (Figure 3.13). Volcano-glacial (hyaloclastite) and glacial (tillite) formations reached up to 50% by volume of the lava piles (Sæmundsson 1996). The recording of the Tjörnes Peninsula (Figure 3.14) began with volcanic activity highlighted by basaltic flows, dated at 9.9 My, 8.6 My and 4.3 My, respectively (Aronson and Sæmundsson 1975), which alternated with the essentially glaciomarine deposits of the first glaciations. These deposits were covered by those formed under temperate conditions: the Tapes zone (4.4–4.0 My) and the Mactra zone (4.0–3.6 My), named after the dominant bivalve shells.

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Figure 3.12. Correlation of glacial deposits identified in Icelandic stratigraphy (completed from Eirikson 2008). In black: normal polarity, white reverse polarity

Figure 3.13. Evolution of Icelandic glaciations during the Plio-Pleistocene. Completed from (Geirsdottir and Eiriksson 1994). It is not impossible, however, that towards 3 My, the extension of the ice cap was greater than that recommended by these authors, promoting rift jump by glacial overload. Notice the rifting-induced shift of Reykjavik location

Cenozoic Evolution of Iceland and the Cryosphere

Figure 3.14. Tjörnes glacial and glaciomarine deposits (source of the log: Verhoeven et al. 2011), with dropstones and shell beds deformed by slumping (© Brigitte Van Vliet-Lanoë)

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3.2.4.2. The Middle Pleistocene Transition (MPT) from 1 My to 0.53 My Throughout peripheral Iceland, the first Pleistocene transition is recorded at about 2.6 My, as in the entire peri-Atlantic zone; thereafter, amplification occurred in successive stages at 2.2–2.1 My, 1.6 My, and then after 1 My (Figure 3.4). Evidence of large glaciations have been recorded since 1.2 My. In the northeastern part of the Skagi Peninsula, glaciomarine sediments located on a coastal or paleo-strandflat platform were sealed by basalts of this age. Significant active volcanism was observed in the northeast rift, from 1.2 to 1 My inland and along the Snaefellness peninsula (Ljósufjöll volcano). Also at this time, the first tuyas forming south of Húsavík (Litla Saltvík) were truncated by temperate-based glaciers (see below Figure 3.30). After this first step, the climate cooled down. The record of glaciations evolved in Iceland with the Middle Pleistocene Transition or MPT (Figure 3.15), well recorded northeast of present-day Vatnajökull. It coincides in particular with MIS 22 (920–880 ky), a very important glaciation, probably still on a temperate basis, responsible for a major discontinuity that marked the beginning of very large glaciations. It also coincided with the formation of the present Snæfells volcano (842 ± 10 ky) (Hardarson 1993). Similar deposits exist in southeastern Iceland, at the foot of the Oræfajökull: the ages of the Svínafell lacustrine deposits range between 1.945 and 0.781 My (Helgason and Duncan 2013). It is also the time of the establishment of the great volcanoes of the south coast such as Eyjafjallajökull (at about 780 ky) or Katla. A major effusive event occurred between 750 and 700 ky in the Reykjanes peninsula (Efstadafjall tuya 717 ± 47 ky; 800 m of ice), which implies an already thick cap. The glacial permanence of Vatnajökull was recorded only 790 ky ago (Helgason and Duncan 2013). In Northeastern Iceland, most of the hyaloclastite ridges (brecciated subglacial hyaloclastite lavas) were emplaced since the MPT, during very large glaciations and especially at about 750–700 ky. Ancient lava flows blocked the flow of the Jökuldalur glacial valley, incised in the middle of the Jökuldalheiði plateau, northeast of Vatnajökull. This event allowed the formation of a very large paleolake, the Múlalón (Guillou et al. 2010; Van Vliet-Lanoë et al. 2010), recorded by 80 m of sandy glaciolacustrine deposits, the Laugarvalla sediments (Figure 3.16). Series A of Laugarvalla overlies the lower basalt dated at 740 ± 27 ky and corresponds to the deglaciation of MIS 19 (790–761 ky). The formation of basal hyaloclastites probably dates from MIS 22 (920–880 ky).

Figure 3.15. Climate change, recorded by δ180 in sea cores (Liesecki and Raymo 2005), releaseof debris in marine sediments (IRD, in red; based on (Helmke et al. 2005; Hodell et al. 2008)) and the Middle Pleistocene Transition (MPT). Incision of valleys in Northeast Iceland (Háslón). Numbers correspond to marine isotopic stages

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Figure 3.16(a). Recording of the MPT within the deposits of the Múlalón paleolake. (Hálslón Dam) (móberg = tuya). Location: see figures 3.17 and 3.30 (source: B. Van Vliet-Lanoë, H. Guillou and A. Guðmundsson, unpublished)

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Figure 3.16(b). Recording of the MPT within the deposits of the Múlalón paleolake (Hálslón Dam) (móberg = tuya). Location: see figure 3.17. Dating: Helgason & Duncan 2003 (source: B. Van Vliet-Lanoë, H. Guillou and A. Guðmundsson, unpublished)

The main deposits constitute the sediments of Laugarvalla B; they are younger than 750 ky and record two glacial events that correspond fairly well to the two stages of MIS 18 (761–676 ky). This period was generally temperate and characterized by limited glaciation (Lisiecki and Raymo 2005). The Laugarvalla C series was deposited during the beginning of the MIS 16 glaciation (676–621 ky). This major glaciation, as well as the later MIS 12 glaciation, also appears to be responsible for a major glaciation on the Yermak Plateau north of Svalbard (Flowers 1997), with a marked increase in IRD accumulation rates in the North Atlantic (Helmke et al. 2005; Hodell et al. 2008). This MIS 16 glaciation is responsible for a major erosive discontinuity and has completely reshaped the Jökulheidi Plateau. The deglaciation of MIS 16 (Termination VII) was the first stage of the Middle Quaternary following the MPT, when the 41 ky orbital cyclicity became dominated by the 100 ky one. The later interglacials MIS 15 and 13 were poorly recorded, as was the glacial episode of MIS 14 (Lisiecki and Raymo 2005; Lang and Wolf 2011). Significant effusive activity was observed in the terminal phase in the Borgaviki valley (640 ± 34 ky) in the north of the island, as well as on the Snæfells volcano (651 ± 14 ky) (Hardarson 1993). The last regional lava flows took place on the surface of the Jökulheidi plateau. Dated between 560–480 ky (Usagawa et al. 1999), they were contemporaneous with

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those of Fnjóskadalur (Akureyri; 546 ± 31 ky (Guillou et al. 2010)) and Snæfells (521 ± 14 ky) (Hardarson 1993). This volcanic activity attests to a significant glaciostatic discharge at the end of MIS 16 and a very long isostatic and magmatic readjustment (section 2.1). 3.2.4.3. Isotopic stage 12, the last very large glaciation (MIS 12) and MIS 10 Little information is available on land in Iceland, but in Northwestern Europe this glaciation was by far the most powerful of the Middle and Upper Quaternary. It corresponds to the Elsterian or Anglian of Europe and the Pre-Illinois B of North America. Intensification of glacial conditions was indicated offshore by a marked increase in IRD accumulation rates in the North Atlantic (Figure 3.4) (Helmke et al. 2005; Hodell et al. 2008) and by a weakening of the thermohaline circulation (Hodell et al. 2008). The maximum glacial extent for Iceland seemed also to occur during MIS 12 (478–424 ky). The main evidence is the formation of tuyas at the beginning or at the end of the glaciation, and especially in a peripheral position with respect to the island. This was the case for Snæfellsnes (Vatnafell 414 ± 11 ky, 150 m, 550 m of ice) or the Reykjanes Peninsula (Borgarfjall, 410 ± 47 ky, 250 m, 600 m of ice; (Guillou et al. 2010)). It was also the case for the tuyas located further inland such as those of Hágongulón (437 ± 38 and 459 ± 47 ky) and Bláfell (460 ± 14 ky), south of Hvítarvatn, located east of Langjökull (Guillou et al. 2010). The hyaloclastite ridge of Fjallgarður (northeastern Iceland), immediately north of Mörðrudalur, has migrated towards the present rift due to ice load (400 ± 100 ky) (Garcia et al. 2003). The Hekla probably had its first eruption (413 ± 13 ky 40Ar/39Ar age on feldspars) at the boundary of MIS 11, a very long interglacial, climatologically similar to the present-day one (Figure 3.15), but started its activity probably at the end of MIS 12. It is likely that the Torfajökull was also formed at this time, and that the Látra morainic arc at the western edge of the Iceland Shelf also corresponded to this major glacial extent and its maximum IRD peak. In any case, the ice cap reached the edge of the platform in Southeastern Iceland, disturbing the flow of North Atlantic Deep Water (NADW; Viviane Bout-Roumazeille, UST Lille, personal communication). During the following glaciation, the Saalian I or MIS 10 (343–374 ky), very large tuyas formed in the interior of the island, such as the Laufafell (373 ± 10 ky, 450-800 m of ice) and the Torfajökull (384 ± 20 ky), or the Kerlingafjöll (336–345 ky) southwest of the Hofsjökull (Guillou et al. 2010; Flude et al. 2010). Many of these tuyas are rhyolitic in composition.

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During the last 500 ky, as a result of these extensive glaciations, the rate of erosion increased to 50–175 cm. kyr-1 (Geirsdottir et al. 2007). In fact, the incision of large valleys occurred mainly after the very large glaciations (Figure 3.15), probably forced by a major glacioisostatic rebound. The subsequent interglacial, MIS 9 (343–280 ky), was relatively warm with three high sea levels. 3.2.4.4. The Saalian II (MIS 8 and 6) The activity of the Fjallgarður hyaloclastite ridge (northeast) persisted during the large glaciations of the Middle and Upper Pleistocene. The Saalian IIa or MIS 8 (280–240 ky) was not a significant glaciation on a global scale, according to the isotope curve of δ18O by Liesecki and Raymo (2005) presented in Figure 3.15. Nevertheless, significant tuyas formed at about 250 ky, particularly the eastern Snaefell and the Heirðubreiðartögl (256 ± 66 ky, 500—800 m of ice) (Guillou et al. 2010). The ice cap gradually thickened, as shown by the morphology of the tuya with several stabilization plateaus. Torfajökull also experienced subglacial activity 255 ± 20 and 236 ± 7 ky (Clay et al. 2015). It was a brief glaciation, not very extensive in Europe, but very cold. It was possible that the formation of these tuyas and tindars occurred just at the end of the 290 ky interstadial, which also corresponded to volcanic activity in the German Eifel and the French Massif Central and a major seismic crisis in Western Europe and the Mediterranean basin. MIS 7 was a complex interglacial, fresh or often considered as a succession of interstadials, but no clear evidence has been observed in Iceland, due to the persistence of the ice caps, probably slightly larger than today. A few ice caps were preserved at that time in Scandinavia (Mangerud 2004). The Saalian IIb or MIS 6 (180–135 ky) was characterized by a mostly temperate-based cap (Van Vliet-Lanoë et al. 2005). In Iceland, this dynamic explains the erosional capability of the ice sheet glacial outlets, which have largely shaped the present morphology (Geirsdottir et al. 2007), barely splayed by the deposits of the last glaciation. From a climatic point of view, MIS 6 was a snip less cold than MIS 2 (Last Glacial Maximum (LGM), 20 ky) on the North Atlantic margins (Mangerud 2004). Isotopic reconstructions of sea level variations showed a significant cooling at about 163 ky (Elderfield et al. 2012), associated with the first Saalian II morainal system, the Drenthe stage in Northwestern Europe. This was the last maximum glacier extent in both Europe and North America (Ehlers et al. 2011): MIS 6c is associated with the storage of considerable volumes of ice at the periphery of the North Atlantic and especially in Greenland (Ehlers 1983; Mangerud et al. 1998; Svendsen et al. 2004). This maximum extension was related to the persistence of fairly mild conditions in the Atlantic Ocean (Funnell 1995). A first thermal minimum (ocean surface temperature) was then reached before about 157 ky

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(Eynaud et al. 2007). The climate then became arid at the end of this stadial, corresponding to the first and largest MIS 6 dust peak observed in the Vostok and EPICA ice cores in Antarctica (Figure 3.17). However, the ocean remained cold with winter sea ice descending far south during the next interstadial (MIS 6b: 157–150 ky) (Margari et al. 2014). This was a major episode of IRD production in the North Atlantic (Figure 3.4), which also coincided with a maximum discharge from the Fenno-Scandian ice sheet via the English Channel (Toucanne et al. 2009).

Figure 3.17. Paleoclimatic record of the last 400 ky in Vostok, Antarctica (modified from Petit et al. 1999)

In Iceland, the intraglacial deglaciation of MIS 6b was equivalent to that of the Bølling (14.5 ky BP; see section 1.3.6), both in the south and in the north of the island. It seriously thinned the cap from 155 ky (Figure 3.18(A)), as shown by the hyaloclastite ridge west of Kleifarvatn (Reykjanes) and the subaerial lava flows of the Jökullheidi near the Hálslón dam. The Kerlingarfjöll volcano was also very active at about 150 ky, with emission of jökulhlaups and formation of juxtaglacial lakes with icebergs. Volcanic activity also occurred at about 150–155 ky in Eyjafjörður (north of Akureyri) (Guillou et al. 2010, 2019). This activity bore witness to a significant and relatively rapid interstadial deglaciation episode.

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At the foot of the Hekla volcano, a subglacial lake formed, in association with jökulhlaups and a collapse of the slope of a parasitic cone. The interaction between the orbital forcing and the millennium-scale variations allows us to divide this lacustrine formation into: an ancient glacial lake, MIS 6c; an interstadiary lake, MIS 6b; and a final stadiary, MIS 6a (Van Vliet-Lanoë et al , 2018). From marine cores, MIS6 b was relatively boreal and humid with very large variations in climate (Margari et al. 2014), similarly to the beginning of the last glaciation during MIS 5c and 4 (110–40 ky). Large interstadials under orbital forcing still existed, notably at about 175 ky. This cold but very wet period in Iceland was related to one of the maximum extensions of the Icelandic ice cap, which probably covered most of the shelf, as in MIS 12. MIS 6a stadial advance ended this glaciation with a context as cold as MIS 2 from the last glaciation. This period corresponded to the thermal minimum of this glacial episode ca 140 ky ago, with a rapid re-englaciation (advance of the Warthe River in Northwestern Europe) which reached Northern Germany, a little south of the Weichselian LGM boundary. The MIS 6 ice cap was thick and extensive and also responsible for ice streams with the first and major glacio-isostatic rebound during the Eemian in South Iceland. The glacio-isostatic load on Iceland at the end of MIS 6 also explains the importance of subglacial hyaloclastites, which were emplaced during Termination II, the Late-Saalian deglaciation. 3.2.4.5. The last interglacial (Eemian) and Termination II Marine Termination II corresponds to a period of simple and rapid deglaciation between 132 ky and 128 ky (Broecker and Henderson 1998), although some authors start it at about 136 ky. Global dust fluxes ceased, and atmospheric CO2 levels increased due to plant recolonization and ocean mass warming. The major European deglaciation was occurring, although in Iceland some subglacial volcanoes still seemed to be developing at high altitudes such as the Prestahnjúkur (Langjökull) (Clay et al. 2015). In the North Atlantic, this state of deglaciation was reached at about 129 ky based on marine boreholes and ice cores (Govin et al. 2015). This early age was confirmed by the deglaciation of the Snaefells Peninsula, which occurred between 135 and 129 ky according to the ages obtained on the subaerial basalts from Ljósüfjöll (Guillou et al. 2010). Significant lava emissions and phreatomagmatism occurred during this deglaciation in the Northern Volcanic Zone and in the south of the island; they were buried beneath interglacial sedimentary deposits. Sand deposits formed during deglaciation and MIS 5e were mainly derived from hyaloclastites emplaced during the previous glaciation (Van Vliet-Lanoë et al. 2007, 2018). Jökulhaups were common during the deglaciation in the north and south.

Figure 3.18. Iceland during the Middle Saalian and Eemian. (A) Location of sites showing strong deglaciation during MIS 6b (about 10 ky). Sites dated in red, probable sites in pink. (B) Extension of Eemian facies (132–113 ky). Red dots: dated volcanic activity and active ridges; in orange: recognized sedimentary deposits (B. Van Vliet-Lanoë and H. Guillou, unpublished)

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The sedimentary record of this interglacial is widely represented in the Northern Volcanic Zone. Locally, glaciolacustrine formations appeared first, then were very quickly overlain by estuarine tidal formations in the coastal zone, outcropping today up to 250 m above sea level, as a result of the glacio-isostatic rebound. Then, emersive lacustrine deposits formed along the Jökulsá á Fjöllum, in the upper Jökuldalur, with an interstratification of fluvioglacial, fluviolacustrine and tillite facies. Estuarine series are also observed in the Vopnafjörður valley to the north-east and, as paleo-estuaries, in the Rangá and Þjórsá valleys to the south, as well as on the strandflat of Snæfellsnes (Figure 3.18(B)). First described in the north (Van Vliet-Lanoë et al. 2001, 2007), the deposits of this interglacial are known as the “Rangá formation” throughout southern Iceland (Van Vliet-Lanoë et al. 2018) (Figures 3.18 and 3.19). This is the first continuous and complete terrestrial record of the Eemian interglacial in Iceland and probably in most of the North Atlantic margins. In the south, this formation overlies discontinuously the tillites of MIS 6b and 6a. Its emplacement was linked to a rapid deglaciation event followed by two marine transgressions marked by the development of coastal mudflats and separated by a complex phase of regression with the development of paleosoils and loess deposits. Thanks to tephrostratigraphy, a better analysis of the interconnections between relative sea level, intensity of regional and marine glaciation in the northern hemisphere or extension of the ice caps is possible during this interglacial. This estuarine infill, with its high eustatic levels and an interstratified regression, records the distal signature of a complex glacial advance well identified in Northern Iceland (Hálslón) (Van Vliet-Lanoë et al. 2010) between 120 and 116 ky BP. Paleo-jökulhlaups, basaltic lava flows and tephra from the volcanoes Hekla, Krafla, Barðarbunga, Askja and especially Grímsvötn have affected the peri-icelandic (offshore) as well as the terrestrial sedimentary record. This sedimentation was particularly significant during the beginning of deglaciation, in association with numerous earthquakes and volcanic eruptions, both in the South Icelandic Seismic Zone (SISZ) and in the northeast rift along the Jökulsá á Fjöllum. During the thermal optimum of the Eemian (127 ky), the end of the first marine transgression south of the island was marked by a Plinian rhyolitic eruption from the Hekla. The lacustrine deposits to the north and south recorded the deposition of a complex basaltic tephra (3 eruptions) from the Grímsvötn volcano during the early thermal optimum of the interglacial (127 ky BP) (Van Vliet-Lanoë et al. 2018); this phreatomagmatic tephra is known in marine cores as the “5e Bas low IV”. It was associated ashore with significant jökulhlaup deposits, implying the persistence of glaciers in west-central Vatnajökull and most probably in Langjökull. The

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stratigraphic position of 5e Bas low IV tephra during the Eemian interglacial is identical to that of the Holocene tephra of Saksunarvatn. In the north of the island, the Grímsvötn caldera also brought along the Jökulsá á Fjöllum valley an early and very significant amount of silt derived from hyaloclastites. This was also the time when the Krafla caldera was formed (Sæmundsson 1991). Askja developed an eruptive activity during Early Eemian as significant as that of the Early Holocene (tephra Askja S, see section 1.3.2.1).

Figure 3.19. Eemian deposits

COMMENTARY OF FIGURE 3.19. – (A) Loess of the Eemian Rangá formation in the south of Iceland, interstratified with basaltic tephra coming mainly from the Grímsvötn volcano. (B) Eemian lacustrine deposits of Kotahæd (Jökulsá á Fjöllum valley, north) reworking hyaloclastites from the Grímsvötn volcano (light gray), two contemporary Grímsvötn tephras from the “5e Bas low IV” of the sea cores (see overview in Davies et al. 2014) and a tephra from the Bárðarbunga (© Brigitte Van Vliet-Lanoë). The cooling leading to the last glaciation can also be traced by the formation and altitude of the tuyas. It was very early and intra-Eemian south of the Langjökull (118.7 ± 17 ky for the Holdufell), but limited to the plateaus, with a thick but cold-base ice cap. The ice sheet was less thick in the north, in relation with a marked extension of sea ice. This was recorded by the amplitude of the second transgression: 165 m in the south versus 120 m in the north of the island. This intra-Eemian glacial advance was attributed to Eemian cooling events (Greenland Stadial GS 26) or to marine cooling events C25–C26 (Van Vliet-Lanoë et al. 2010, 2018). This corresponds fairly well with what was observed in Southern and Eastern

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Greenland, with a marked glacial advance at 119 and 117 ky (Fronval et al. 1998; Irvali et al. 2016).

Figure 3.20. Reconstructed apparent eustatic curve in Southern Iceland from Rangá formation data (Van Vliet-Lanoë et al. 2018) compared to NGRIP (2004) (purple) and Medina-Elizalde’s eustatic curve, 2013 (green; from intertropical corals)

A second significant deglaciation event then occurred, again during the Eemian, in relation to the warming of the GI 25 in Greenland, a large interstadial linked to a resumption of the well known thermohaline circulation in the northeast Atlantic, itself controlled by the delayed maximum insolation in the southern hemisphere. It was marked by the arrival of tropical waters at 50° N between 120 and 116 ky and the persistence of a mildness over the eastern North Atlantic from 116 to 113 ky (Fronval et al. 1998; Rasmussen et al. 2003). On the other hand, Greenland and the Northwest Atlantic were gradually cooling simultaneously.

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This late deglaciation was recorded in Iceland by a second marine transgression in the south (Figure 3.20) between 116 and 113 ky BP, in association with a major melting of the Icelandic ice cap of the same order as that recorded for the Holocene optimum in the Eyjafjörður sector (disappearance of the eastern Vatnajökull ice cap) (Striberger et al. 2012). The ignimbritic flows of the caldera of the Krafla volcano indicate that at least part of the Northern Volcanic Zone was still ice free (116–110 ky) (Sæmundsson 1992). 3.3. The last glacial episode and its deglaciation 3.3.1. The Weichselian 3.3.1.1. The transition with the last glacial episode A major cooling occurred at 113 ky (end of Eemian, in Greenland GS 25, C 24 of cores). The ice sheets reformed where rainfall was still abundant, in Southern and Western Iceland as stressed by Broecker and Denton (1990) (section 3.1), in relation to the Irminger Current, then thickened and finally began to spread lately during interstadials. The glaciers were cold-base in altitude, with the permafrost coming down on the plateaus below 500 m altitude (Van Vliet-Lanoë et al. 2010). 3.3.1.2. The beginning of the Weichselian glacial During the beginning of the Weichselian glacial, Iceland was most probably covered by an ice sheet partially covering the southern marine platform (Petursey tuya, 113 ± 5 ky). The base of the central Hengill volcano is considered to be of the same age and reflects a succession of stadials and interstadials (Sæmundsson 1995). In the central depression of southern Iceland, the arrival of glacial tongues was posterior to 110 ky: the Langjökull developed in altitude further south (Figure 3.21) and immediately north, as did the Kerlingarfjöll (Clay et al. 2015). More than 550 m of ice covered the top of the Torfa caldera (Owen et al. 2012). In Reykjanes peninsula, the Skálamælifell tuya is 40Ar/39Ar dated at 94.1 ± 7.8 ky (Jicha et al. 2011). Glaciers developed a little later in the north: the Bruarjökull, north of the Vatnajökull, extended to the north of the Hálslón dam at about 103 ky. Dating of the rhyolites of the Krafla volcano yielded ages between 85 and 90 ky (Sæmundsson et al. 2000), indicating a limited and highly variable englaciation of the Northern Volcanic Zone. Subaerial lava flows were still emplaced at 80 ky along the Fallgarður ridge in the Jökulsá á Fjóllum valley (Guillou et al. 2010).

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Figure 3.21. Modalities of the extension of the Icelandic ice sheet based on the heights and ages of the Langjökull subglacial volcanoes and dated facies in sea cores from the northern and northwestern shelves. Diagram taking into account glacio-isostatic subsidence (Van Vliet-Lanoë, unpublished)

From MIS 4 (< 75 ky BP), the formed ice sheet most likely reached the present coast with quite considerable glaciation to the south and west, and overflowed onto the shelf to the south. It corresponded to the glacial advance of Karmoy in Norway (Mangerud 2004). The local ice caps were thick (500–800 m), but not very extensive. This is normal in a context of high precipitation in the south and still “mild” temperatures at low altitude. The north, already arid, was drying out due to the extent of sea ice. This is probably the time when rock glaciers were being formed on the coast and at high altitude due to aridity, such as those covering marine deposits that were probably Eemian (> 43 cal ky: age limit of 14C) (Guðmundsson 2000b).

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At the beginning of the Middle Pleniglacial interstage (MIS 3), the Torfajökull subaerial ignimbrites described at Þórsmork (behind the Eyjafjällajökull) towards 55.3 ky attest to limited deglaciation as in the southern hemisphere at about 55–60 ky (Bazin et al. 2013; Veres et al. 2013; Guillou et al. 2019). This period corresponds to the Bø interstadial in Norway. Rock glaciers covering marine deposits continued to evolve on the north coast. Subsequently, a significant indirect index of ice load was the formation of numerous subglacial volcanoes within a larger ice sheet at ca 50–40 ky BP in Central Iceland (Guillou et al. 2010). This recording, via the altitude of their plateaus, indicates an ice thickness of 500–800 m. This episode was also recognized in Norway (Mangerud 2004) and in Western Europe (including France) (Van Vliet-Lanoë et al. 1999). It generally corresponded to the maximum advance of glaciers on the eastern margins of the Atlantic. 3.3.1.3. The Middle Pleistocene interstadial A glacial retreat was recorded between 30–40 ky (Kristjansson and Guðmundsson 1980; Eirìkson et al. 1997). Peaks of sedimentary discharge (IRD, dropstones) were observed in the northwestern peninsula in Djúpal (Geirsdottir et al. 2002) at 34, 32 and 26–30 ky. A retreat of the ice cap was recorded in the Reykjanes Peninsula at ca 35 ky, in association with an apparent modest sea level rise (+5 m) associated with a coastal paleoridge (Sæmundsson and Norðdahl 2002). This transgression was contemporaneous with the Ålesund interstadial in Norway (Larsen and Mangerurd 1989) and the marine transgression of the Middle Weichselian reconstructed by Svendsen et al. (2004) for the Scandinavian Arctic. 3.3.2. The Last Glacial Maximum 3.3.2.1. Background The Last Glacial Maximum (LGM) sensu lato covers the period between 26 and 18 ky. It corresponds to the coldest period and the lowest sea level. Nevertheless, the LGM period is currently set from 23 cal ky to 18 cal ky (Clark and Mix 2002). However, the precise timing of the LGM and the maximum extension of the Icelandic cap are poorly constrained, although the deglaciation that followed is reasonably well known. Several authors have postulated the existence of an extension of the ice sheet that would have reached the shelf edge during the LGM (Norðdahl and Halfidason 1992; Ingólfsson et al. 1997; Hubbard et al. 2006; Patton et al. 2017). Their hypothesis is based on the elevation of the undated tuyas,

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the potential age of the last Latra Moraine on the western shelf of Iceland, and the terminal moraines observed by multibeam prospecting, but not dated, at the edge of the Icelandic shelf (Spagnollo et al. 2009; Patton et al. 2017). It was also based on the occurrence of a tephra related to the classic Vedde Ash formation (Katla volcano, 11.8 cal ky BP), the tephra of Skógar, that was considered as the tracer of Younger Dryas (Norðdahl and Halfidason 1992), but was further proved to be the result of several eruptions. L. Lane and her collaborators have demonstrated that there are at least three successive rhyolitic tephras from Katla, dated at 15–13.5 and 11.8 ky respectively (Lane et al. 2012), two of which have been identified on land in Iceland and Scandinavia. 3.3.2.2. Observations In Norway, the LGM was centered on 25 ky. Western Norway recovered regularly high precipitation via the recurrent reactivation of the North Atlantic Drift, in relation to the Ålesund interstadial (35–27 ky). This suggests that the maximum glaciation should be shifted later in Iceland, and more specifically in the north of the island due to the winter extension of sea ice and probably to the importance of the East Greenland and Icelandic cold drifts. During the LGM and especially during its deglaciation, the importance of ice streams on the Greenland and Scandinavian Atlantic margins (Ottesen et al. 2005, 2009) or along the margins of the Arctic Ocean (Bjarnadóttir et al. 2014) led to the formation of floating masses of glacier ice and local grounded terminal moraines down to the edge of the Icelandic shelf. Undated traces of large icebergs deformed the Icelandic shelf and shelf edge (multibeam imagery) (Patton et al. 2017). A significant glacial retreat of the South Icelandic ice sheet was recorded onshore during the LGM, in the Reykjanes Peninsula, at ca 25.5 ky cal BP and 20–22.5 ky (Eiriksson et al. 1997; Johanesson et al. 1997), in association with a still high relative sea level. IRD records of Garðar Ridge cores show recurrent calving episodes since 65 ky, indicating the unstable nature of the ice sheet since at least the Heinrich 6 event (widespread ice break-up) in relation to temperate DO events. This is corroborated by the age of the Þórsmörk ignimbrite (53 ky). Only the southern coasts and the shelf were really occupied by a spreaded ice sheet like the Edgeøya one in Eastern Spitsbergen, under the thermal impact of the Irminger current and thermohaline circulation (AMO). In the southeast, the Vatnajökull probably emitted ice streams at the level of the Breiðamerkurjökull glacier (Jökulsárlón), the Skeiðarárjökull and the Skafta Valley during the LGM period, as well as along the Markarfljὀt (behind the

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Eyjafjallajökull). To the west, north and east, the outflow from the main ice sheet could thus form floating tongues of ice (Figure 3.22) spreading or advancing away from the outlet of the present glacial valleys (Van Vliet Lanoë et al. 2005, 2010; Principato et al. 2016). Large ice-free areas (nunataks) (Ingólfsson 1991; Guðmundsson 2000b) appear to have persisted during the LGM along the coastal mountains, particularly in Northwestern, Northern and Eastern Iceland.

Figure 3.22. Recording of ice sheet behavior, subglacial and subaerial eruptive activity during the Last Glacial Maximum based on the dating of tephra recognized in marine and glacial cores (A), subglacial volcanoes (B) and debris discharges in the seas surrounding Iceland (simplified from Van Vliet-Lanoë et al. 2018)

This restricted glaciation pattern is consistent with the absence of offshore deposits in the Eyjafjörður (Akureyri Fjord, north; Figure 3.23), from 27 cal ky to

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18 cal ky (Andrews et al. 2000, 2003). The Eyjafjörður is disconnected from the main ice sheet due to the altitude of the mountains bordering it to the south, but constitutes a watershed in its own right given the relief that surrounds it. Several morainic formations occupy the fjord and deglaciation is complete during Younger Dryas (true Vedde Ash resting on a perched marine beach) (Van Vliet-Lanoë et al. 2005).

Figure 3.23. Extension of the ice cap during the LGM

COMMENT ON FIGURE 3.23. – Morainic arcs towards 300 m depth (according to Hafro.is; (Spagnolo and Clark 2009; Patton et al. 2017)) and marine sedimentary fans (according to (Vogt et al. 1980)). The deep moraines (at about – 200 m) are in thick black lines, the others in thin lines. White: true ice cap. Light blue: detectable ice streams (morphology). Light grey: probable extent of ice shelves to the LGM . Red bar: no ice passage during the Weichselian. A: Reykjavík; B: Borgarfjörður; A: Akureyri; H: Húsavík; E: Egilsstaðir.

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The northwestern peninsula appears to have had an independent ice cap, with calving glaciers emitted from a watershed located near the center of the peninsula (Principato et al. 2006). These glaciers were bordered by block fields, mature periglacial landforms and a significant number of rock glaciers (Guðmundsson 2000b), indicating relative aridity. Each nunatak was characterized by its own glacial system, independent of the ice sheet (Hoppe 1982). Offshore in the Djúpáll basin (northwestern peninsula), ice growth was recorded at 21–22 ky BP (Andrews et al. 2000, 2002b; Geirsdottir et al. 2002), followed rapidly by a retreat on the shelf at 20 and 21 cal ky BP (Principato et al. 2006). Sonar imagery (HAFRO, GEBCO), also used by Spagnolo and Clark (2009) and Patton et al. (2017), shows an ice cap edge recorded on the bottom by a multiple series of end moraines (Figure 3.23). 3.3.2.3. Ice streams The compilation of the work of several authors (Vogt et al. 1980; Thors 1982; Sturkell et al. 1992) make it possible to delineate several polyphase submarine sedimentary deposits of glacial origin along the edge of the Icelandic shelf (Figure 3.23). Marine seismic profiles show 2 km thick polygenic sedimentary accumulations in the outlet basin of Öxarfjörður and Skjálfandi (Húsavík Bay) (Thors 1982; Sturkell et al. 1992). The most significant sedimentary deposition areas are located downstream of the large valleys occupied by ice streams, whose flow patterns converge (Figure 3.23). Their size and life span are controlled by the size of the watersheds upstream of these valleys and their supply in ice by the Vatnajökull, Langjökull and, to a lesser extent, the Hofsjökull. The most obvious traces of rapid flow are concentrated in Húnafjörður and Húnaflói (Figures 3.23 and 3.24). The same is true for Hunafloadjúp and Skagafjörður, which are mostly fed by the Langjökull. Imagery data from Spagnolo and Clark (2009) show the junction of the two ice streams on the offshore platform, with lateral moraines. Skagafjörður shows clear traces of a rapid flow. The Látra moraine is exposed at the end of the western platform (Figure 3.23) at a depth of 200–350 m. Its height is 20–30 m, but it includes some reliefs that can reach 100 m (Ólafsdóttir 1975). The seismic profiles attest to its polygenic character. This area is located between the great glacial sedimentary sprawl and disconnected from the supply by the glacial valleys of the platform. The Látra moraine most probably corresponds to an ice extent older than the Weichselian, anchored, dynamic and of short duration as described by Dodeswell et al. (2014) across from Disko (West Greenland). Also the altered till and striations observed on Grímsey Island (Figure 3.23) (Einarsson 1967; Norðdahl 1991), 40 km north of Husavik, may have been emplaced during an earlier and more extensive glaciation, given their degree of weathering. To establish a bottom moraine (adherent glacier), the ice sheet must be thick, with a relatively temperate base and a very low sea level,

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which corresponds mainly to MIS 16, MIS 12 and MIS 6 with a global sea level ≥ – 150 m (Figure 3.21).

Figure 3.24. Ice rivers. Terminal moraine arcs in the Eyjafjörður (Greinivík, Spot 3 image) and ice streams tracks upstream of Húnafjörður (Spot 4 IR image) with the 9.3 and 10.3 ky BP moraines identified in yellow

The last phase of sedimentation on the seaward side next to the Látra moraine comprises laminated sedimentary facies under floating ice cover, younger than 35.6 cal ky BP but older than 18.6 cal ky (Syvitski et al. 1999; Andrews et al. 2000). This would correlate well with the 22–21 ky period observed in Djúpáll basin (Andrews et al. 2000; Geirsdottir et al. 2002), which corresponds perfectly with a floating tongue or ice shelf emitted during a glacial surge. Given an apparent sea level of 0 m at 30 ky (global – 50 m), it is possible to estimate that the relative sea level during the LGM was about – 70 m (global – 120 m). This results in a minimum thickness of anchored glacier ice in the order of 130–150 m. The classic thickness of an Antarctic ice shelf is about 250 m, but it can float from 50 m thickness upwards.

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3.3.3. Deglaciation and the Holocene 3.3.3.1. A two-step deglaciation The deglaciation of the North Atlantic margins is under the control of the evolution of ocean circulation and a southern heat transfer by surface waters. The relatively warm Atlantic waters of the Irminger Current reached the outer shelf of northern Iceland at 18.6–16.3 ky cal BP. The thermomohaline circulation reactivated northwards about 14.7 ky cal BP ago. The Bølling interstadial (Termination Ia; Figure 3.10) was particularly warm in spring and summer due to orbital forcing and was associated with a very rapid sea level rise, the Meltwaterpulse 1a, reaching 15 m in 340 years, destabilizing the ice shelves or the rim of the ice sheet (Ingolfsson et al. 2010). The western North American ice sheet disappeared completely at this time (Grégoire et al. 2016), and the summer sea ice cover also disappeared completely as far as the Fram Strait (Müller et al. 2009; Stranne et al. 2014). The chronology of coastal glacial retreat in Northwest Iceland is the best studied, and is found to be earlier (from 15–17 ky) (Andrés et al. 2019; Hodell et al. 2019) than the periods of rapid retreat of ice fronts in East Greenland (13.0–11.5 ky). Nevertheless, the deglaciation of Greenland was not homogeneous, as it was later in the south (11–10 ky) and especially in the west (10.5-7.0 ky) in relation with the cold east Greenland Drift (Sinclair et al. 2016). In Iceland, the Húnaflói ice shelf disappeared at ca 15 cal ky (Syvitski et al. 1999; Andrews et al. 2000; Eiríksson et al. 2000; Geirsdóttir et al. 2002). Iceberg scours and evidences for significant calving were observed northeast of Skagi to Malland with consistent 36Cl datings (Andrés et al. 2019). Subaerial lavas from the Þeistareykir shield volcano and Krafla caldera in the north of the island showed that at least parts of the Northern Volcanic Zone deglaciated rapidly during Bølling/Allerød (Sæmundsson 1973, 1991). Bølling deglaciation appeared to be major, as Vedde ash (11.8 cal ky) was observed immediately north of Dettifoss in association with Askja tephra (Preboreal 11.5 ky) (Van Vliet-Lanoë et al. 2020), while Askja S tephra (11.0 cal ky BP) was found in a sandur only 20 km north of Vatnajökull (Van Vliet-Lanoë et al. 2010b). In the south, considerable deglaciation also developed from the Bølling (Geirsdóttir et al. 1997; Jennings et al. 2000, 2007; Principato et al. 2006; Licciardi et al. 2007): the Jökuldjúp basin (west of Reykjanes) was ice-free from the beginning of this interstadial. The Reykjanes Peninsula also became permanently ice free from 14.5–15 ky cal BP (Jóhannesson et al. 1997). Sea ice disappeared in the north of Iceland. On land, it was the time when the discharge from the slopes and the seismicity associated with deglaciation triggered large landslides, which were reworked by moraines and especially by rock glaciers at the end of the Late Glacial and Preboreal periods (Figure 3.25).

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Figure 3.25. Altitude of the tuyas (section from Vatnajökull to Öxarfjodur and Langjökull to the coast) in relation to the volume of the Icelandic cap (modified and reinterpreted from Walker 1965; Licciardi et al. 2007). Data in red: (Van Vliet-Lanoë et al. 2018)

COMMENT ON FIGURE 3.25. – All these tuyas are supposed to have formed during the last glaciation; the available K-Ar or tephra ages are in red and show inherited landforms. The Langjökull at 104 ky (early glaciation) is in orange. The 10Be ages are in fact synchronous with stages of limited top de-icing or re-icing. Most of the ice sheet deglaciation occurred during the Bølling (Van Vliet-Lanoë, unpublished). Subsequently, the climate cooled from the end of the Bølling, following a series of events including the emptying of the Canadian Agazziz Lake (Murton et al. 2010), minimal solar activity and a volcanic paroxysm, and sea ice re-extent (Müller et al. 2009). This period is marked by a series of glacial advances of increasing intensity, the Bølling cold event (BCE), the Allerød cold event (ACE) and finally the Younger Dryas event, characterized by two successive advances, classically interpreted as leading to the formation of an ice sheet reaching the coast via outlet glaciers. However, it is unlikely that a real ice sheet could form in 2,000 years, given the aridity of the time (extensive sea ice). The situation was much more

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complex. The moraine of the Fjallgarðar Ridge plateau (east of Mörðrudalur) shows a 36Cl age close to 13.2 ky, indicating a very limited advance of the ice sheet at the end of the Bølling (Bölling Cold Event) (Meriaux et al. 2013). The outlet glaciers of the residual ice sheet flowed down again at the end of the Bølling interstadial and reached the coast such as at Öxarfjörður (section of Röndin, near Kópasker). No real ice shelves were formed anymore, only calving tidal-front ice streams, as evidenced by the occurrence of iceberg scouring tracks at Húnaflói until the Younger Dryas (Eiriksson et al. 2000). In the Skjálfandi and Vopnafjörður valleys, the terminal moraines of the outlets issued from the Younger Dryas relict ice cap were emplaced back inland or along the coast such as at Borgarfjörður (Figure 3.26). The very large cobble beaches show evidence of a sustained distal calving at Malland, northeast of the Skagi Peninsula, corroborating the 36Cl dating (Tanarro et al. 2019). Mapping observations, Vedde tephra extent and cosmogenic datings exclude the reconstruction of an ice sheet during the Younger Dryas in the northwest, north and northeast of the island due to aridity (Sæmundsson et al. 2012; Andrés et al. 2019; Tanarro et al. 2019; Van Vliet-Lanoë et al. 2020b). Subsequently, in Northern and Northeastern Iceland, the ice cap no longer reached the coast. This is corroborated by the limited importance of glacio-isostatic rebound in the north, reinterpreted from data by Rundgren et al. (1987). Northeast of Reykjavík, the termination of an outlet glacier issued from the Langjökull cap was marked by deltaic glacial front deposits associated with marine shorelines (Ingólfsson et al. 1995). This was also the case at Borgarfjörður, where a surge from this ice cap spreaded towards Hvalfjörður in the south and calved on the strandflat further north. This pattern shows that Iceland was not exceptionally cold during the Late Glacial final episode, and rather suggests an intermediate glacier behavior between those of Norway and Greenland, with, as in Norway, a sensitivity to glacial surges exacerbated by the major role of a regional mild current, that of Irminger. 3.3.3.2. The Holocene evolution of glaciers The beginning of the Holocene was still very unstable in the North Atlantic (Werner et al. 2013), with long periods of temperate water inflows into the Arctic basin, separated by brief cooling events associated with an extension of sea ice. These temperate water inflows were linked to a powerful North Atlantic drift along the Svalbard coast, then through the Fram Strait into the Arctic Basin. Generally speaking, the world’s glaciers regress, or even disappear with brief re-advances during the cooling episodes of the Pre-Boreal. The Laurentide and West Greenland ice sheets regressed after the Younger Dryas.

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Figure 3.26. Sedimentary facies linked to deglaciation

COMMENT ON FIGURE 3.26. – (A) Late Glacial lateral moraine of the Borgarfjörður glacier tongue, washed by marine transgression (Termination Ib; Águst Guðmundsson, Hafnafjörður). (B) Glacial blocks released by the same, partly floating, glacial tongue to the northwest of Borgarfjörður on glaciomarine sediments. (C) Traces of melting of an iceberg with a sediment-bearing base in a 5 m high coastal sandy aggradation, north of Hella (Southern Iceland) (photos B, C ©Brigitte Van Vliet-Lanoë).

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The last ice sheet to disappear was the Labrador one. It survived under optimum thermal conditions until about 6 ky BP, due to the importance of the East Greenland oceanic drift. In Iceland, the majority of the ice sheets and ice caps, which had already regressed during the Bølling, retracted mainly during the Preboreal (Figure 3.27). Glaciers responded to changes in ocean circulation and sea level rise as seen during the Late Glacial and the Eemian interglacial. Mild water flowed towards the northwest during the resumption of the Irminger Current and the apparent eustatic rise (Figure 3.10) of the Early Holocene destabilized glacial outflows reaching the coast. Deglaciation continued, less and less interrupted by glacial advances (Van Vliet-Lanoë et al. 2020b). A limited glacial advance in southern Iceland was observed at the beginning of the Preboreal at about 11.2 cal ky, when the Búði moraines were formed, after the maximum flooding linked to the glacio-isostatic rebound. Various smaller glacial advances were also recognized, notably at 10.3 and 9.3 cal ky BP.

Figure 3.27. Chronology of deglaciation in Iceland between 12 and 8 ky (Van Vliet-Lanoë et al. 2020b)

The second glacial pulsation occurred at 10.3–10.2 ky, still in the same logic as the previous advances, corresponding to the Erdalen events in Norway. In the north of Iceland, these were still rapid advances of polythermal glaciers (cold temperature

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at the surface, close to zero at depth) on a permafrosted ground. In the south, the glacial tongues were spread out and probably already with a water-lubricated and a temperate 0°C basal temperature. It was also the period of the major eruptions of the Grímsvötn volcano (Saksunarvatn tephra). The ice rapidly receded afterwards, the ocean temperature reaching values similar to the current temperature and the air temperature rising more than 2°C above the current values. The glacial advances at 9.3 ky correspond to the majority of the ancient terminal moraines still observable downstream from the maximum glacial advance of the Little Ice Age. They were erratic in flow direction with respect to the previous glacial system (Kaldal and Vikingsson 1990). This advance has also been recorded, notably in the northwestern peninsula (Drangajökull) (Ingolfsson et al. 2016) and west of Vatnajökull, before being covered by lavas from the Þjórsá (8.6 ky). It most probably corresponded to the last presence of permafrost at medium altitude in Iceland before the Holocene climatic optimum. After a last, very limited advance at 8.2 cal ky BP, glaciers disappeared east of Vatnajökull (Striberger et al. 2012), and no moraine evidence is currently visible. The only advance recorded indirectly by sedimentology and 10Be is that of the Langjökull ice cap, due to the contribution of precipitation by the Irminger current (Figures 3.25 and 3.27). This cooling, recorded between 8.5 and 8 ky cal BP, is associated with a re-extension of sea ice to the north of the island; it is probably correlated with an emptying event of Lake Agassiz, sustained volcanic activity (Cooper et al. 2018) and a restriction of the Irminger Current, limiting precipitation. The Greenland ice sheet also retracted, but resisted up to 7 ky to the southwest, just like the Canadian Ungava ice cap, still chilled by a powerful and cold East Greenlandic drift. The latter only disappeared at 6 ky. During the Holocene thermal optimum, temperatures were about 3°C higher than at the present time, such as in Northern Quebec at 57° N. Svalbard saw its average temperature rise by 4°C compared to the current one; the boreal forest even reached the northern coast of Europe, and the oak the south of the White Sea. The permafrost shrunk and disappeared from Western Europe. Nevertheless, some Icelandic glaciers seemed to remain on the Vatnajökull and Langjökull (Bjornsson 2017) and Icelandic permafrost subsisted above 1,000 m. The system stabilized at ca 5.2 ky, when the peri-arctic continental shelves were delayed in inondation due to their merging resulting from isostatic rebound (Greenland and North America); their maximal flooding was reached at 6 ky. From 4.4 ky, the glaciers and then the ice caps were reconstituted in Scandinavia as well as in Iceland. The correlative extent of the sea ice led to the formation of deep polar Arctic waters (ADPW). Then, from 3 ky, the layer of cold and brackish superficial waters from the sea ice-melt remained at the surface of the Arctic basin.

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This was the Neoglacial period, controlled by the progressive decrease in global insolation (Figure 3.28) since 8 ky, a decrease to which were superimposed marked cooling episodes, forced by a decrease in solar activity, particularly that of the Iron Age (2.8 ky), and periods of high volcanic activity. The progressive cooling of the East Greenland current since 5 ky was interrupted by a brief warming between 3.9 and 3.5 ky, before its resumption from 3.5 to 2 ky.

Figure 3.28. Holocene evolution of glaciers in the North Atlantic area. Note the longer time for glaciers restoral along the Norwegian coast. Modified and completed according to (Balacio et al. 2015; Gjerde et al. 2016). The dotted line corresponds approximately to the equilibrium line of permanent snow

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A relatively mild period (300–500 AD) then led to the arrival of Irish monks (the Papar) on the island. Then, the glaciers thickened again, especially during the very wet and cold period (500–750 AD) which preceded the Medieval Optimum (800-1300 AD), a period which saw the arrival of Norwegian settlers, in association with a warming of the waters of the North Atlantic drift. Finally, the strongest glacier advance occurred during the Little Ice Age, especially since 1650 (Spörer and Maunder solar minima). Most of the retreating moraines visible at the periphery of the ice caps correspond to the very last advance (1880), the older advances (1650) being often transformed into rocky glaciers (Figure 3.29(C)). Permafrost re-extended in altitude above 600 m (Van Vliet-Lanoë and Guðmundsson 2020). With the extension of sea ice, precipitation was reduced in the north, and the glaciers of the Tröllaskagi Plateau (Figure 3.29(A)) also evolved into rock glaciers.

Figure 3.29. Current rock glacier and glacier morphologies

COMMENT ON FIGURE 3.29. – (A) Cirque rock glaciers, drifting from glaciers, Tröllaskagi. (B) Rock glacier derived from a landslide induced by deglaciation seismicity; Skagi (photos A, B © Águst Guðmundsson, Hafnafjördur). (C) Rock glacier derived from a moraine of the 17th century, west of the Bárðarbunga (photo LMIs). (D) Múlajökull (Hofsjökull)1, Little Ice Age deglaciation. 1 https: //notendur.hi.is/oi/icelandic_glaciers.html.

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As the climate warmed, precipitation became more abundant, ice shelves broke up, and large rivers flowing into the Arctic poured cooler, fresher and lighter water onto the surface of the Arctic Ocean. This phenomenon allowed a rapid re-extent of the sea ice, but slowed down the thermohaline circulation and its thermal input. On the other hand, this global increase in precipitation induced abundant snowfalls in altitude on perennial ice sheets such as that of Greenland, causing their thickening and an increase in their albedo, even if their coastal zones tended to melt as in the years 2000–2006 or 2017-2019. Nevertheless, the retreat of calving glaciers in Southwest Greenland has been extremely limited during the last 9.1 ky BP (Dyke et al. 2017). Currently, Icelandic glaciers are melting very fast, especially in the south of the island, in connection with high precipitation related to the oscillation of thermohaline circulation (or AMO) and a slightly warmer Irminger Current. For an ice cap like the Vatnajökull, whose snow is polluted with basaltic volcanic ashes (mainly from the Grímsvötn volcano), the significant decrease of the albedo from the ash load results in an acceleration of the melting and the reappearance of jökulhlaups. The glaciers could reduce quite considerably, or even perhaps disappear temporarily like at the end of the Eemian, and the volcanic activity could increase correlatively. Nevertheless, the present-day permafrost on Tröllaskagi seems to resist this warming in altitude, above 1,000 m (section 3.4.4.1). This situation is reminiscent of what happened during the Eemian in Iceland, Greenland and to the North Atlantic drift, with a second warming at the end of the interglacial. It results from the 3,000-year-long gap between the interglacial thermal optimum of the northern hemisphere and that of the southern hemisphere. The warming of the Southern Ocean, the largest and most extensive in the intertropical zone, was witnessed around the year 1000 AD by thermohaline circulation. Like during the Eemian (cold event GI 26), we have just experienced a cooling episode, linked in particular to the solar activity, better known as the Little Ice Age (1300 to 1820 AD), followed by a rapid warming such as there was at the end of this GI 26 Pleistocene interglacial. This natural evolution of the Holocene interglacial (Van Vliet-Lanoë 2018) suggests above all a thermal control by the ocean, like during the Bond events (or the Dansgaard–Oeschger events during the Ice Age), accentuated by the Atlantic Multidecadal Oscillation (AMO), even if human impact probably plays on this natural background. Nevertheless, the duration and amplitude of the Holocene orbital forcing are less intense than those of the Eemian. We are currently entering a period of solar cycles that are getting longer since solar cycle 24, which prejudges the end of the solar component of the current warming.

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3.3.3.3. Neogene and Quaternary erosion Iceland has been exposed to a boreal climate since the Middle Miocene. Then, since 9 My, it has been cyclically glaciated by temperate-based mountain glaciers, especially since 5 My. This is evidenced by the very marked glacial erosion relief affecting the Northwestern Peninsula, Flatey and Tröllaskagi in the north, and the relief of the east coast. The latter were also partially flooded by the sea as far as the upstream cirques of the Northwestern Peninsula, due to thermal and/or gravitational subsidence (Figure 1.8 in Volume 1). As discussed in Chapter 1 of Volume 1, the strandflat surrounding the island is derived from a Neogene marine abrasion surface, which has been entrenched by glaciers since 9 My. The dating of the lava flow sealing the lower platform (+ 5 m) of Skagi yielded a K-Ar age of 1.8 My, and the platform capped by the 137 ky old lavas of Gerðuberg (Snæfellnes) is perched at + 40 m. These paleosurfaces were thus maintained by recurrent abrasion. Given the early glacio-isostatic rejuvenation at the coast, these surfaces are exposed, during the very beginning of interglacials and during interstadials or during low relative sea levels (– 70 m during the LGM), to the abrasive and clearing action of coastal ice and sea ice (Lisitzin 2002; Are et al. 2008). This sea ice is still forming episodically at present. This means that the seabed extending the strandflat offshore has been systematically stripped of all deposits (except glacial outflows basins) by glacial dynamics since 9 My. The first real ice caps only appeared during the Gelasian (ca 2.7 My ago) (Geirsdottir and Eirıksson 1994), as suggested by the emplacement of the first subglacial eroded volcanoes (tuyas) south of Húsavík (Litla Saltvik). Valley incisions related to isostatic rebound under mechanical erosion control appeared at about 3.3 My and 2 My, apparently associated with an episode of uplift of the Atlantic margins (Japsen et al. 2014), then at 1.2 My, 750 ky, 350 ka and 150 ky, rather related to major deglaciation episodes (Figure 3.15). Hyaloclastite ridges are more or less correlated with long temporary deglaciations (≈ 10 to 20 ky) of very extensive ice sheets on the surface and over time (≈ 100 ky) (Van Vliet-Lanoë et al. 2018). The Fjállgarðar hyaloclastite ridge in the NVZ was mostly functional at 0.70 and 0.40 My, corresponding to the end of two very large glaciations (MIS 16 and MIS 12; figures 3.15 and 3.16). The significance of erosion increased after 500 ky, during the Middle Pleistocene Transition (Geirsdóttir et al. 2007) and the transition towards glacial cyclicity at 100 ky. The large active canyons (Jökulsá á Fjóllum and Jökulsá á Dal) were last active

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at the end of MIS 6, but their major incision is much older; it probably dates minimally from MIS 12 (400 ky), as evidenced by the presence of lithified tillites. Denudation rates in the unglaciated and interglacial boreal periods were of the order of 5 cm/ky, with correlative deposits in the shelf channels of the order of 6 cm/ky (Geirsdóttir et al. 2007).

Figure 3.30. Examples of ancient deposits. (A) Paleotuya truncated by very old tillites south of Húsavík (seen towards the north, cliff height: 35 m). (B) Lithified glaciolacustrine deposits in the bottom of the Jökulsá canyon at Fjöllum (12 m above the present canyon bottom) (© Brigitte Van Vliet-Lanoë)

Glacial erosion has been especially important since 500 ky, during the deglaciation periods, when glaciers of very large volume became polythermal or

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temperate based, with erosion rates between 100-300 cm/ky. This led, for the last 500,000 years, to an erosion of about 100,000 km3 for an erupted magma volume of 40,000 km3. The erosion generated by glacially dependent episodes was therefore much greater than the magmatic inputs. Volcanic inputs remained very localized during the last 500 ky, more particularly at the level of rifts, as attested by the radiometric or stratigraphic dating of the Quaternary volcanic deposits. These volcanic emissions were always typical of an active hotspot, probably impacted by the whole isostatic discharge resulting from mechanical erosion. 3.4. Iceland today, its climate and vegetation 3.4.1. The climate The Icelandic climate is oceanic subarctic and, despite its latitude, strongly moderated by the influence of the Gulf Stream to the south and the intermittent presence of the Irminger Current to the west of the island. Snow is abundant in winter, especially above 400 m. The south is very humid and rainy (Figure 3.31), under the influence of the positive NAO and especially positive AMO, as has been the case since 1980. Maximum precipitation occurs in the southern part of Vatnajökull. The northeast of the island is much colder and drier, in relation to the East Icelandic current resulting from the East Greenlandic drift, and, depending on the circumstances, the extension of the sea ice. Rainfall increases with altitude, thus explaining the position of the ice caps over the most powerful reliefs of the island: the Bárðarbunga, the Grímsvötn (Vatnajökull), the Katla (Myrdaljökull), the Hofsjökull and the Langjökull. The wind is often strong depending on the position of the Icelandic cyclonic depression which generally deepens south-east of the southern edge of Greenland. Normally located in the southwest of the island, the cyclonic depression is responsible for winds from the north or northeast over the Central Zone of Iceland, accentuating the aridity and erosion of the soils, and violent easterly winds along the south coast of the island. Westerly winds are rare. Very strong storms are frequent, especially in late summer or fall, often associated with metric coastal storm surges. In fact, the polar jet present at altitude undulates according to the negative character of the NAO or stabilizes in periods of positive NAO (years 1994–2004). In periods of climatic cooling, such as during the years 1960–1987, the jet goes back down in altitude and pushes the depressions southward.

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Figure 3.31. Average annual rain and temperatures for the period 1931–1960. (1) < 600 mm; (2) 600–1,199 mm; (3) 1,200–1,999 mm; (4) 2,000–3,999 mm; (5) > 4,000 mm. Annual average temperature, Veðurstofa Íslands2

3.4.2. Ocean circulation and climate 3.4.2.1. Marine currents The surface ocean circulation in the North Atlantic is controlled by winds, forming a relatively mild (≥ 8°C in the south) and salty current, the North Atlantic Drift (NAD), extending the Gulf Stream northward. A cooler branch (2 to 5°C), the Norwegian Current, enters the Barents Sea and propagates to Novaya Zemlya, and the main branch runs along the west coast of Spitsbergen to enter the Arctic Basin. In the southwest of Iceland, a branch breaks off and rises towards the north of the island, forming the mild Irminger Current (4–6°C), characterized by relatively salty water (34.9–35.0 psu). This current is unstable and can deflect southward in periods of climate cooling. This current is very important, as we have seen in this chapter (section 3.21). Its expansion northwest of Iceland is a major factor in deglaciation. A coastal current flows clockwise over the Icelandic basaltic shelf and is characterized by low salinity near the coast in spring and summer, when runoff is at its maximum and mixing in the ocean at its minimum. It is an ancillary component of the Irminger Current.

2 https://notendur.hi.is/oi/climate_in_iceland.htm.

Figure 3.32. (A) Map of North Atlantic marine currents. Fine black arrows: cold deep salty water (NADW). (B) North Atlantic circulation in 1883 (reconstructed from (Mohn 1887)) and 1943 (US Army chart) NAD: North Atlantic Drift; IC Irminger Current, EGD: East Greeland Drift (Van Vliet-Lanoë 2018)

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In the north of Iceland, it is called the North Icelandic Irminger Current; in winter, the katabatic winds coming down from the Greenlandic ice cap push this current back southward. Its average speed is 5 to 10 cm.s-1. During the warm episode of 1940–1945, the secondary branch of the current circled the island from the west, whereas during the cold years of 1880–1883, it was on the contrary the East Icelandic drift that circled Iceland from the east (Figure 3.32(B)). The marine cores in the Icelandic Sea are thus a very good recorder of the evolution of the climate. The thermal input derived from the Gulf Stream and relayed by the Irminger current is a major tracer of warming episodes related to intertropical warming, whether for reasons of orbital forcing, solar activity or El Nino events. This is crucial for the south and west coasts of Iceland, the north and east coasts remaining under the dominant influence of cold water from the East Greenland drift. This phenomenon is very sensitive to the level of deglaciation via surface temperatures north of the Denmark Strait (intrusion of the Irminger current north of Iceland), leading to the disappearance of the winter pack ice in the Icelandic Sea and the retreat of glaciers. On the other hand, excessive inflows of cold melt water, for example during the expulsion of the Arctic pack ice forced by the winds, pushes it back to the south of the island. 3.4.2.2. The sea ice around Iceland and the climate Sea ice forms on the ocean surface during the polar winter when the seawater temperature drops below –1.8°C. At the end of winter, the ice can be as thick as 1.5 to 2 meters along Northern Greenland, not counting the snow that accumulates. Its buoyancy is due to the difference in density between ice and seawater. Sea ice acts as an insulator between the ocean and the atmosphere. Because of its presence, it limits evaporation from the ocean and thus the oceanic caloric input and precipitation, at least for the winter period. When it is covered with snow, the albedo of the sea ice is almost identical to that of fresh snow. On the other hand, the sea ice emits infrared radiation and promotes regionally a significant cooling of the overlying atmosphere: the ice-free ocean absorbs 90% of the incident energy, whereas the sea ice without snow already reflects 70% of it. Its extent will significantly modify the radiation during cold periods. As a layer of seawater cools to its freezing point, its density increases and it sinks to the bottom, bringing less salty water to the surface. The entire column of water of the same concentration will have to cool down to reach freezing point. This column is usually about 10 meters thick, and is separated from the dense, salty water by a

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halocline with a strong salinity gradient. The sea ice then forms frasil, that is a slurry of ice crystals on the surface, which will gradually consolidate. Sea ice contains even more salts excluded by the crystallization of the ice the faster the speed of ice setting and the rougher the sea. The migration of the brines excluded from the crystallization takes place at the opposite side of the freezing zone and causes in the long term a depletion in salts of the old ice floes. During its formation, by expelling cold brine, the sea ice modifies the temperature and density of the underlying salty water which plunges towards the ocean floor, leading in the long run to a desalination of the colder surface water, retroactively promoting the extension of the pack ice. It therefore plays an important role in thermohaline oceanic circulation and for the climate of the planet; it is also a factor in the maintenance of the Greenland ice sheet. After the Holocene thermal optimum, it will reform at about 5,500 years cal BP (Figure 3.33).

Figure 3.33. Holocene history of the Icelandic sea ice. (A) Holocene extension of the sea ice on the North Icelandic Shelf (Cabedo-Sanz et al. 2016). Gray bars: cold phases. (B) Index of the recent extension of the sea ice around Iceland (NSDIC)

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The average winter sea ice cover is generally located in the northwest of Iceland, about 100 kilometers from the coast. Its existence was first mentioned by the Irish monk Dicuil in 795 AD. It can reach the coast of the north-western peninsula, the northern and eastern coast of the island during cold periods such as those of the Little Ice Age, the springs of 1882, 1902 and 1911 or the Ice Years 1965–1971 (Figures 3.33(B) and 3.34), allowing the random arrival of polar bears on the Northwest Peninsula. Its maximum extension, also encompassing the south of the island, was noted in 1881. Such an extension, also noted in 1917 (Figure 3.34), was penalizing for the fishing industry and especially cod reproduction. During the Little Ice Age it was located between the British Isles and Iceland. On the other hand, the sea ice may temporarily disappear from the North Atlantic during episodes of varying length. Such a situation was observed in 1930, and also occurred during the interstage under the control of a solar excursion, the beginning of the Bølling (14,500 years cal BP) (Andrews et al. 2009), at the end of the LGM in the Barents Sea, or during the Eemian thermal maximum (127 ky).

Figure 3.34. Maximum extension of sea ice in April 1917. It reached Reykjavík in 1911. The dashed line marks the southernmost sea ice limit, and the small triangles the limit of drifting ice in Greenland (source: Danish Meteorological Institute)

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3.4.3. Soil, people and climate 3.4.3.1. Humans, fauna, flora and climate in Iceland Scandinavia and Greenland benefit from an arctic fauna adapted to the climatic conditions, with a greater diversity in the south, despite a certain endemism induced by glaciations. Iceland being an isolated land, like Spitsbergen, the fauna is very poor apart from birds, marine mammals and some great travelers like the polar fox and polar bear, drifting with the pack ice, as was the case during the Little Ice Age or during glacial periods. Irish and Scottish monks, the Papar, colonized southern Iceland as early as the 5th century, via Orkney, Shetland and the Faroe Islands. This is evidenced by religious books and objects found in Iceland, and by crosses carved in caves dug into interglacial sediments (Figure 3.35). The Vikings, Scandinavian warriors and merchants, attacked, colonized, and explored the Atlantic coasts between 300 AD and 800 AD, during a cold period responsible for starvation in Scandinavia and Western Europe. They arrived locally in Iceland as early as the 7th century. On the other hand, the main phase of colonization of the island took place mainly between 870 and 930 AD, corresponding to a relatively temperate period from a climatic point of view, with a mild climate linked to a relatively positive NAO and subarctic winds allowing westward navigation. This was essentially the founding of a nation or Landnáma (874 AD) by an agricultural population. The Vikings also brought with them slaves and cattle from the British Islands. At the onset of the Medieval Thermal Optimum, in 985 AD, they arrived in Southern Greenland, then covered with willow and birch forests, where Erik the Red founded a colony that disappeared in the course of the 15th and 16th centuries, during the cooling of the Little Ice Age and the expansion of the Thule Eskimo culture. The other mammals currently present on the island accompanied the Vikings and even the Papar: these are the domestic animals and rodents that arrived with the hay reserves for the cattle. The various species currently raised in Iceland were almost all imported from Norway or brought back from plunder during the colonization in the 9th century. These are mostly small species such as goats and sheep, mainly of the Nordic varieties. Pigs are close to the Scottish and German varieties and are very present in the local toponymy. This is also true for dogs, close to Scandinavian dogs and Scottish shepherds. The Icelandic cow was imported in the same way from Norway. The historical varieties have been well maintained and adapted to the harsh climate. When the Vikings arrived in numbers at about 874–930 AD, they could only bring two horses per drakkar. So they selected the best and toughest from the herds of Norway and

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Germany, but also from the ponies of Northern Ireland, Scotland and the Shetlands. At the end of the 10th century, imports ceased and the Icelandic horse evolved in an endemic way, mixing the various breeds and adapting to the harshness of the climate.

Figure 3.35. Evolution of Icelandic habitat

COMMENT ON FIGURE 3.35. – (A) Papar huts generally associated with cellars excavated in the Eemian formation (preserved below topping lava, Ytri Rangá Valley). (B) Reconstructed Viking long-house (11th century, destroyed by the eruption of Hekla 1; Ytri Rangá valley). (C) Sænautasel farm (1843 AD) occupied for a century (Fjallgarður ridge). (D) Note the thermal insulation by the loessic peat (© Brigitte Van Vliet-Lanoë). During the medieval climatic optimum, Icelandic breeders selected their horses according to criteria of coat color, small size, robustness and hardiness, the horse being satisfied with lean food such as lichens, shrubs and seaweed. Their presence was reported in the Landnámabók.Viking agricultural production was mainly food, fodder, with some grain production (barley, rye, spelt) and rare vegetables (turnips, cabbage). The absence of predators promoted cattle and pig breeding. That of sheep

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intensified in the century following the Landnáma. This production was supplemented by the collection of eggs and hunting (geese, ducks, swans, partridges, puffins, seals). The techniques of conservation were varied (drying, smoking, elaborated dairy products: butter, skyr). Reindeer and mink are of recent importation (after 1945). 3.4.3.2. Icelandic vegetation Given the youth of Iceland, it is not possible to go back beyond 16 My, the maximum age of the currently emerged layers. Plant colonization of the island probably took place through more or less discontinuous landbridges, especially during low sea levels (tectonic and glacio-isostatic regression of the Late Oligocene, at about 23 My, and very low level of the base of the Tortonian, at 11.6 My) (Denk et al. 2005; Dickson et al. 2011; Wappler and Grímsson 2016). Plant biodiversity was highest in the Middle Miocene (15 My, see Table 3.1). Oligocene Flora of Turgaï 25 My Pinacea: Metasequoia, Ginkgo, Magnolia

Middle Miocene Boreal or arcto-tertiary flora, 15 My (Icelandic data)

Lower Pliocene 5 My (Icelandic data)

Pinacea: Picea, Pinus, Abies, Pinacea: Picea, Pinus, Abies, Larix, Tsuga, Gingko, Larix Sassafra, Magnolia

Alnus, Betula, Populus, Fagus, Salix

Alnus, Betula, Populus, Fagus, Salix

Alnus, Betula, Populus, Fagus, Salix

Quercus, Acer, Platanus, Thuja, Taxodium, Petrocarya, Juglans, Ulmus, Corylus, Vitis, Viburnum, Rhus

Quercus, Acer, Platanus, Thuja, Taxodium, Petrocarya, Juglans, Ulmus, Corylus, Vitis, Fraxinus, Viburnum, Rhus, Carpinus, Rhododendron, Ephedra

Sorbus, Rhododendron, Potamogeton

Table 3.1. Evolution of Icelandic flora during the Neogene (based on Meyen 1987; Wappler and Grimsson 2016)

The Pliocene subsidence of the Atlantic margins and that of the residual continental plates, linked to the opening of the Atlantic, progressively isolated the island, even during periods of low sea levels. The island was forested during the Upper Miocene (15–5 My), with a flora characteristic of a warm temperate environment (Metasequoia, Magnolia, Sassafras, Pterocarya). Beech seems to have been abundant (Fagus sp.), indicating a humid climate under the influence of the future Irminger Current. From the Pliocene onwards, before the regular recurrence of the Pleistocene glaciations, the boreal forests dominated by Pinus, Picea, Abies,

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Larix, Betula and Alnus attested to a cooling of the climate, even at low altitudes.The tundra became differentiated about 4 My ago. Present-day Icelandic vegetation is impoverished and mostly of Palearctic type. About 20% of the flora survived the last glaciation via nunataks emerging from the ice sheet. The poverty in species can be attributed to the isolation of the island rather than to the climate. Most species are common with Norway (97% common species), the northern British Isles, but much less so with Greenland (66%) and the rest of the Arctic. Some seeds come from strandings in coastal areas, especially those transported by the Gulf Stream and the Irminger Current. This explains the low contribution of Greenland to plant colonization. The first plant to germinate on the volcanic island of Surtsey emerged in 1963 was a tomato plant! Icelandic flora is characterized by 470 endemic and naturalized vascular species, including 37 species of mosses, 7 species of horsetail, 10 species of ferns including 5 varieties of the primitive Botrychium (Figure 3.36(B)), about 290 dicotyledons, including the famous Dryas octopetala (Figure 3.36(D)), and numerous species of ericaceous plants, 145 monocotyledons and 2 orchids. It should be noted that a significant number of the Graminae and Papilionaceae is of northern European affinity and was brought with fodder by the early Viking settlers. The presence of ferns is linked to the snowy character of the Icelandic climate, maintained by the Irminger current, which allows frost-free winters at the ground level for the boreal varieties. Many species of native grasses reproduce directly by sprouts instead of seeds to shorten the vegetative cycle: this is the case of viviparous sheep’s-fescue (Figure 3.36(A)). Dwarfism and prostration (creeping vegetation), such as the polar willow (Figure 3.36(E)), are a phenotypical adaptation to cold. At the beginning of the Holocene (Preboreal), postglacial recolonization was very limited due to a frequently cold and dry context; only peat bogs began to develop in wet oases. Plant colonization was carried out via nunataks and floating seeds. Between 9,000 and 8,500 years cal BP, pioneer species such as juniper and birch (Betulus pubescens) developed. During the Holocene climatic optimum (8,000 years cal BP), the mild temperatures and heavy rainfall promoted the birch forest, which reached its greatest extension at about 6,000 years cal BP, as in Eyjafjörður. The “forest” was probably to rise to about 600 m, but not beyond, the domain of prostrated or discrete tundra (figures 3.37(A), 3.37(B)). Then, with the gradual cooling associated with Holocene orbital forcing, the birch regressed in altitude for the benefit of peat bogs and wet moorlands, as in Northern Ireland and in Scandinavia. This cooling became more pronounced at about 5,500 years cal BP, the period of Bronze Age cooling recorded all around the North Atlantic, in association with very high storminess and sea ice.

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Figure 3.36. Icelandic flora. (A) Viviparous sheep’s fescue. (B) Botrychium fern. (C) Uniflower cottongrass. (D) Eight petal mountain-avens (Dryas octopetala). (E) Polar willow. (F) Lichens and crowberry (© Brigitte Van Vliet-Lanoë)

Before the arrival of the Vikings in the 9th century, 30% of the island was still covered with birch forests, with dwarf birch, willow and juniper bushes, mainly at less than 400 m altitude (Figure 3.37(C)). A cold and arid desert currently occupies the central plateau (Figure 3.37(A)), volcanic massifs and mountainous areas such as the Tröllaskagi Peninsula. It consists of glaciers (about 10% of Icelandic territory), areas of bare volcanic rock

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(about 23%), open tundra (about 13%) and alluvial or glacial sands (about 3%). Wet depressions allow the formation of peaty sedge oases, often occupied by palsa or lithalsa above 600 m (Figure 3.39).

Figure 3.37. Quasi-natural plant associations in Iceland

COMMENT ON FIGURE 3.37. – (A) Prostrated tundra (600 m). (B) Shrub tundra (600 m; willows and dwarf birch). (C) Clear forest of pubescent birch (400 m). (D) Dense birch forest (200 m) surmounted by an altitudinal gradient of vegetation. (E) Cryptogamic crust. (F) Dense cover of moss over lava flow (Racomitrium lanuginosum) (© Brigitte Van Vliet-Lanoë).

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Very often, the “bare” soil is covered by a cryptogamic crust consisting of lichens (Figure 3.36(F)), including Cetraria islandica (Iceland moss), filamentous algae and mosses, forming a kind of natural geotextile that protects against runoff and deflation. Basaltic flows more than 200 years old are colonized by a thick, gray and woolly dry moss, Racomitrium lanuginosum, reminiscent of early Paleozoic soils (Figure 3.37(F)). 3.4.4. Soils and erosion 3.4.4.1. Permafrost One of the characteristics of Iceland is the presence, still today, of a deep permanently frozen ground, the permafrost, continuous above 800 m altitude with annual average temperatures below –3°C. It is mainly located at high altitudes and in the northern continuations of the central Icelandic plateau such as at Tröllaskagi or along the northern edge of Vatnajökull and other ice caps (Figure 3.38). This situation is fairly comparable to that of Norway because of the mildness conveyed by the Irminger Current and the North Atlantic Drift. One of the characteristics of Icelandic permafrost is its high ice content, which allows its persistence at depth in a significant part of Northern and Central Iceland. This ice content frequently induces the formation of isolated rock glaciers or glacier-derived rock glaciers (Guðmundsson 2000b) (Figure 3.29). Its melting also triggers landslides such as those that occurred on Askja volcano in 2012 and Snæfells volcano in 2018 during hot and rainy summer episodes. Sporadic permafrost also exists in the form of palsal fields above 600 m to the south-central part of the island (TMA – 3°C to – 1°C; Figures 3.38 and 3.39), or rock glaciers to the edge of the northern coastal cliffs. Its thermal situation is critical with global warming, at least on the surface and at medium altitude (Etzelmüller et al. 2007), but the forms at higher altitudes (> 1000 m) survived the thermal optimum and still survive (Tanarro et al. 2019). Another source of permafrost is related to the descent of cold air from the ice caps. Historic moraines (Little Ice Age) in the center of the island are often colonized by rock glaciers (Figure 3.29(C)). Permafrost also remains within the terminal moraines of Eyjabakkajökull in the northeast (Figure 3.39(A)), and at much lower elevations in the south of the island, such as in those of Skeiðarárjökull. Insulating sandy sediments such as the Askja pumices allow the development of open-system hydrolaccoliths (pingos) (Figure 3.39(D)), massive ice domes fed by underground lateral drainage. The formation of palsa or lithalsa fields, or mounds of segregated ice (Figure 3.39(C)), is the most common form of permafrost in the oases of the central Icelandic desert.

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Figure 3.38. Map of current permafrost. Dark blue: continuous; light blue: discontinuous; green: deep seasonal frost; yellow: shallow seasonal frost (< 1 m). D: hydrolaccolith or pingo in open system; P: palsa or lithalsa; S: stringbogs (Van Vliet-Lanoë and Guðmundsson 2020; map based on photo-interpretation)

Figure 3.39. Glacial and periglacial figures associated with permafrost

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COMMENT ON FIGURE 3.39. – (A) Permafrost moraine with imbricated push-ridges (active glaciotectonics; Eyjabakkajökull). (B) Decametric polygons (25 m) of thermal contraction (deep seasonal frost). (C) Fields of degraded palsas in a peat bog at 600 m altitude, (discontinuous permafrost). (D) Hydrolaccoliths in open system (pingos), partly in process of thermal degradation (Askja, continuous permafrost, 800 m altitude) (images A, C, D © LMIs; photo B © Brigitte Van Vliet-Lanoë). Nevertheless, their temperature close to 0°C, especially for the palsas, makes it very sensitive to climate change, especially the thermal impact of the snow cover. At present, the palsa fields are most often in the process of degradation, as in the warm years 1930–1950 (Friedman et al. 1971); they redeveloped during the 1960-1980s, leaving small circular lake fields in the peat bogs. Most of the development of permafrost at mid-altitude during the Holocene bears witness to an initial phase of formation in the Preboreal, followed by disappearance during the Holocene thermal optimum. During deglaciation, the permafrost must have invaded the freshly deglaciated areas, increasing the efficiency of runoff and lateral drainage, especially during the first jökulhlaups (Van Vliet-Lanoë et al. 2020b). The last permafrost re-extension dates from the Little Ice Age, more particularly from the Maunder Solar Minimum (ca 1650–1750 AD) (Van Vliet-Lanoë and Guðmundsson 2020), with traces of decametric polygonal ground up to 250 m in the north and 400 m in the south. Cosmogenic dating of the rock glaciers on Tröllaskagi has shown that they formed as early as the Bølling and survived the Holocene optimum above 1,000 m (Tanarro et al. 2019). Palsas and lithalsas are very sensitive to warming, especially due to the presence of stagnant water nearby. The last period of degradation of the palsas occurred during the warm period of 1930–1950. The 1970s and 1980s saw their re-extension according to Friedman et al. (1971). Currently, permafrost melting, or thermokarst, is noticeable, especially on rocky slopes and in the palsa fields. On bare and windy ground, thermal contraction is always functional, even at low altitudes, due to the possibility of cold waves in winter (most recent: –15 to –19°C and winds of 150 to 250 km/h, 2, 7–8, 10–12 December 2019). 3.4.4.2. Soils Icelandic soils are dominated by basaltic volcanic materials (Arnalds 2015) and are classified as andosols in the World Soil Taxonomy (established by the Food and Agriculture Organization, FAO); they are generally fairly alkaline, fertile and have a fairly high nitrogen content (C/N ≈ 11) due to the presence of cyanobacteria.

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Figure 3.40. Periglacial soils. (A, B) Thufurs (Þeistareykir and Krafla volcanoes); note the deformation of the tephra within the inherited shape of a desiccation network (frost desiccation) (© Brigitte Van Vliet-Lanoë). (C) Corded (striated) peat bogs and thufur fields (left). The white islets correspond to residual morainic massifs (© LMIs)

Soils are characterized by limited coherence (thixotropic at water stauration) and good water retention, which makes them very sensitive to erosion, liquefaction and landslides. The clay fraction is of allophane (amorphous) or imogolite (tubular) types, which crystalized into nontronite (swelling clays) in pre-Holocene formations by alkalinization in cold and arid environments, or into halloysite and kaolinite in acidic geothermal sites. Brown andosols are the most common soils in Iceland, especially on loess or loess-rich soils; they are often enriched in tephra, with early browning during the Holocene (10,000 to 8,000 years). Their most frequent morphology in uncultivated areas is thufur, a hummocky microtopography linked to seasonal surface frost, which forms within 4–5 years (Figure 3.40).

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Figure 3.41. (A, B) Solifluction tongues. (C) Small frost-sorted soils. Shallow seasonal frost (< 1 m), except A (© Brigitte Van Vliet-Lanoë)

Peat soils are Histosols with at least 20% organic matter; these soils are often several meters deep due to eolian accretion and behave as moderately acidic. Other andosols generally have a relative pH close to neutral, unless they are rich in basaltic tephra, which may alkalinize them. Andosols in non-peaty moist depressions show traces of redox and contain less than 12% organic matter; they are often enriched in loess and tephra and are characteristic of areas with palsas (discontinuous permafrost) or corded peatlands (deep seasonal frost) at lower elevations, or even fossils at sea level as in southern Iceland. Vitrisols correspond to skeletal soils, poorly vegetated, under deflation regs and low in organic matter, common in the central desert of the island. It is the domain of large structured periglacial soils. True cryosols are rare and functional soils are associated in altitude with continuous permafrost, especially in the north of the island or on the reliefs north of Reykjavík. They are often associated with sorted periglacial forms (Figure 3.41), mostly inherited from the Late Glacial (areas of nunataks or early deglaciated landforms). Solifluction is common and stretches thufurs, sorted soils and alpine grasslands in slope direction at all altitudes.

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3.4.4.3. Humans and erosion Since colonization, agricultural clearance (mainly by burning), the exploitation of wood for construction or for fuel and the metallurgy of low furnaces as also the overgrazing of sheep have led to the disappearance of these natural woodlands, which today cover only 1.1% of the island’s surface area (0.2% forest and 0.9% bush). Volcanic activity since colonization has also led to a reduction in forest cover. Since 870 AD, Iceland has gone from having a continuous forest cover below 400 m to a degraded forest around 920 AD. Then it evolved towards an anthropogenic herbaceous tundra since the 14th century. In altitude, regs now prevail. Major jökulhlaups (section 1.4) ravaged agricultural land in the Southern Zone of Iceland, accentuating the impact of volcanic eruptions and storms from ca 1000 AD. The increase of storms during the 15th—18th centuries accentuated wind erosion (Greipsson 2012), especially inland, allowing the settlement of now vegetated dune fields (Figure 1.41). Another typically Icelandic form of erosion is rofaborð, a form of residual erosion in the loess related to swirling winds and sheep tracks (Figure 1.40). Currently, the inland desert (Figure 3.42) is the direct result of the influence of erosion on the continental dune zones reworking jökulhlaup deposits, sandurs, tillites and tephras, especially since the Little Ice Age, and exporting them from inland to the sandy coasts or to the offshore. Coastal dune ridges also are highly exposed to wind erosion.

Figure 3.42. Soil erosion map, simplified from the CORINE database (modified from Arnald 2006)

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3.5. References Andrés, N., Palacios, D., Hodell, L.M., Fernández, J.M. (2016). The origin of glacial alpine landscape in Tröllaskagi peninsula (North Iceland). Cuader. Invest. Geogr., 42(2), 341–368. Andrés, N., Palacios, D., Sæmundsson, þ., Brynjólfsson, S., Fernández-Fernández, J.M. (2019). The rapid deglaciation of the Skagafjörður fjord, northern Iceland. Boreas, 48, 92–106. Andrews, J.T. (1970) A geomorphological study of post-glacial uplift with particular reference to Arctic Canada. Inst.British Geogr., Special publications series no. 2. Andrews, J.T., Harðardóttir, J., Helgadóttir, G., Jennings, A.E., Geirsdóttir, Á., Sveinbjörnsdóttir, Á.E., Schoolfield, S., Kristjánsdóttir, G.B., Smith, L.M., Thors, K., Syvitski, J. (2000). The N and W Iceland shelf: Insights into Last Glacial Maximum ice extent and deglaciation based on acoustic stratigraphy and basal radiocarbon AMS dates. Quat. Sci. Rev., 19, 619–631. Andrews, J.T., Hardadottir, J., Stoner, J., Mann, M.E., Kristjansdottir, G.B., Koc, N. (2003). Decadal to millennial-scale periodicities in North Iceland shelf sediments over the last 12,000 cal yr: Long-term North Atlantic oceanographic variability and solar forcing. Earth Planet Sci. Lett., 210(3–4), 453–465. Andrews, J.T., Darby, D.A., Eberl, D.D., Jennings, A.E., Moros, M., Ogilvie, A. (2009). A robust multi-site Holocene history of drift ice off northern Iceland: Implications for North Atlantic climate. Holocene, 19, 71–78. Are, F., Reimnitz, E., Grigoriev, M., Hubberten, H.W., Rachold, V. (2008). The influence of cryogenic processes on the erosional Arctic shoreface. J. Coast. Res., 24(1), 110–121. Arnalds, O. (2006). Iceland. In Soil Erosion in Europe, Boardman, J. and Poesen, J. (eds). John Wiley, London. Arnalds, O. (2015). The Soils of Iceland. Springer, Dordrecht. Aronson, J.L. and Sæmundsson, K. (1975). Relatively old basalts from structurally high areas in central Iceland. Earth Planet. Sci. Lett., 28, 83–97. Balascio, N.L., D’Andrea, W.J., Bradley, R.S. (2015). Glacier response to North Atlantic climate variability during the Holocene. Climate Past, Discuss., 11(3), 2009–2036. Bartoli, G., Sarnthein, M., Weinelt, M. (2006). Late Pliocene millennial-scale climate variability in the northern North Atlantic prior to and after the onset of Northern Hemisphere glaciation. Paleoceanogr. Paleoclimat., 21(4) [Online]. Available at: http://doi.org/10.1029/2005PA001185. Bazin, L., Landais, A., Lemieux-Dudon, B., Toyé, M.H., Veres, D., Parrenin, F., Martinerie, P., Ritz, C., Capron, E., Lipenkov, V.Y., Loutre, M-F., Raynaud, M.C., Fischer, H., Masson-Delmotte, V., Chappellaz, J., Wolff, E.W. (2013). The Antarctic ice core chronology (AICC2012). Pangaea [Online]. Available at: https://doi.org/10.1594/ PANGAEA.824894.

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Conclusion Brigitte VAN VLIET-LANOË, René MAURY and Hervé GUILLOU with the collaboration of Françoise BERGERAT and Laurent GEOFFROY

Iceland is not only a splendid island whose large ice caps top huge active volcanoes and which is crossed by a rift that forms the geological boundary between America and Europe (Figure C.1). Its history is attached to that of the North Atlantic and, although young, it bears witness to a complex deep geodynamic evolution, presented in Volume 1, which goes back to the Paleozoic. The formation of this island is the result of a mainly basaltic volcanism, linked to a deep thermal anomaly with a debated meaning, also discussed in Volume 1, and whose summit part (hotspot) is located beneath the Vatnajökull glacier. The coupled functioning of this hotspot and the Mid-Atlantic Ridge has resulted in the insularization of an oceanic island with a very thick crust (up to 40 km), which has undergone several rift-jumps and, frequently, the glacio-isostatic loading, probably since 9 My but especially since the Upper Pliocene (3.2 My), in association with a major glaciation of the island. It cannot be excluded that the relocations of Neogene and Quaternary rifts were partly controlled by particularly important glacial episodes, such as dated at 6.7 My, 2.6 My and 0.8 My, with the rifts widening under glacio-isostatic loading. Rates of mantle partial melting are the resulting volcanic activity are exacerbated by deglaciation events, including the current global warming.

For color version of the figures in this chapter see www.iste.co.uk/vanvliet/iceland2.zip Iceland Within the Northern Atlantic 2, coordinated by Brigitte VAN VLIET-LANOË. © ISTE Ltd 2021. Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

 

Figure C.1. Geological map of Iceland, originally at 1/600 000e, H.J Jóhannesson (2014). Náttúrufræðistofnun Ísland – Icelandic Institute of Natural History (available at: https://en.ni.is/resources/publications/maps/geological-maps)

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C.1. The Icelandic magmas: emplacement and interpretation The genesis of Iceland is linked to its spectacular volcanism, mainly basaltic, whose production (80,000 km3 of lava per My during the last 16 My) is lower than that of Hawaii but higher than that of other intra-oceanic volcanic islands such as Réunion or Tahiti. Iceland has about 30 major active volcanic centers: numerous shield volcanoes and fissure fields, mainly located in rifts and major fracture zones. Central volcanoes have underlying magma reservoirs either shallow (about 2–3 km deep) and/or deeper (12–20 km). The volcanological singularity of the island lies in the interactions between magmas and glaciers, leading to the establishment of sub-glacial buildings of tuya (Herðubreið) or hyaloclastite type ridges. The compositions of Icelandic basaltic lavas differ from that of the depleted oceanic basalts (MORB-N) of the Mid-Atlantic Ridge by the presence of enrichments coming from mantle components of the plume. The tholeiitic basalts are largely dominant at the level of the large volcanic structures of the rifts. They derive from relatively high partial melting rates of the mantle; it has been shown that these melting rates increase slightly but significantly during decompression linked to deglaciation, which also results in increased magma production. Alkaline basalts, which are much less frequent, are found in an “off-axis” position on the Snæfellsnes Peninsula and the Vestmannaeyjar Islands. Finally, intermediate basalts between the two previous types, called transitional basalts, form the large central volcanoes of the south (Eyjafjallajökull, Hekla) and southeast of the island (Öræfa volcanic chain). Lavas of intermediate composition (icelandites) and fairly abundant rhyolites are associated with the basalts. The latter have isotopic characteristics similar to those of the basalts, and result either from the partial melting of deep metabasalts (rifts), or from the fractional crystallization of alkaline basaltic magmas (Snæfellsnes). The complex geochemical signature of the Icelandic basalts reflects their derivation in varying proportions from a suboceanic depleted upper mantle similar to the source of MORB-N, and enriched mantle components typical of intra-oceanic islands (EM-1, EM-2 and HIMU). The nature of the enriched components found in the sources of these lavas, as well as the high 3He/4He ratios of these lavas, argue for a deep origin of the Icelandic plume (see Chapter 3 of Volume 1). The source of the Öræfajökull transitional lavas contains an original component enriched in 87Sr, 207Pb and 208Pb. This peculiarity has been interpreted either as being linked to ancient terrigenous or pelagic sediments subducted and then incorporated into the deep mantle, or as being due to slivers of recycled oceanic crust in the upper mantle, or finally as reflecting contamination by continental fragments present in the substratum of this volcano. It is currently difficult to decide between these hypotheses given the available geochemical data.

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C.2. Climatic variations during the Ice Ages: the importance of the North Atlantic and the Icelandic threshold for thermohaline circulation Due to its position in the Central Atlantic and its current partial glaciation, Iceland is a unique recorder of climate change. Indeed, its position makes it a key witness of the major changes controlling the superficial geodynamic evolution of the North Atlantic, but also of the evolution of the oceanic circulation and thus of the climate since the Neogene, through the North Atlantic drift and thermohaline circulation. While the glaciation of the southern hemisphere began during the Upper Eocene following the thermal isolation of Antarctica, sea ice was already present at that time (38 My) in the vicinity of a shallow, salty North Atlantic. The first mountain glaciers appeared ca 20 My ago on the Alaska Coast Mountains. It is not excluded that they were also formed since the beginning of the Neogene with the assistance of the nascent Mid-Atlantic rift, notably in Greenland, on the ancient Caledonian suture. Since 40 My, the thermal subsidence of the Atlantic margins and mainly of the Iceland-Faroe Ridge, submerged since the Neogene, gradually increased the depth of the shoal that blocked the overflow by the North Atlantic cold waters – more particularly via the Rockall Basin – whose Jurassic thinned crust was more susceptible to thermal subsidence. These waters from the deep, cold and salty North Atlantic sea ice were able to mix with the rest of the Atlantic from 35 My, as the Panama Isthmus progressively emerged, and gradually gave rise to the thermohaline oceanic circulation as we know it today. This process accentuated global glaciation by facilitating the continental storage of ice on an orbital forcing background modulating the solar energy input. The subsidence of the Greenland–Iceland–Faroe Islands threshold was faster than the global glacio-eustatic decline, as evidenced by the presence of continental lake sediments under the SDRs of the Greenland–Iceland Ridge. The first glaciations of the North Atlantic Zone were recorded in Greenland at 9.6 My with the last evidence of migration of European vegetation to Iceland. The oldest presently outcropping Icelandic basalts date back to 16 My and the first evidence of glaciers preserved in Iceland are synchronous with those of Greenland, that is 9 My old. The sedimentation of the platform has been dominated by glacial contributions since the major world glacial phase known at 6.7 My, at the very beginning of the Messinian. The Tjörnes sedimentary formation in Northern Iceland is the most continuous record in the northern hemisphere of the Early Quaternary Ice Age together with the Skaftafell region in the south of the island. The last important phase of subsidence of the Greenland–Iceland and Iceland–Faroe Ridges occurred during the Upper Pliocene, ca 2.6 My ago, and was synchronous with an episode of

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uplift of the Scandinavian and Greenlandic margins and with the first Quaternary glaciations, characterized by a 40 ky cyclicity. The last revolution occurred during the Middle Pleistocene transition, which involved the occurrence, from 0.8 My, of very large glaciations with a 100 ky cyclicity, and the entrance of small volumes of deep arctic waters via the Fram Strait. The next step will occur when the Fram Basin will be sufficiently open and deep to allow the connection of cold, salty arctic bottom waters with the current thermohaline circulation. In the longer term, it is the migration of the North American plate that will condition the minimum of our ice age in the future. Recording what happened during the last interglacial in Iceland sheds light, at our scale, on the possible sequence of future natural climate events as a direct consequence of orbital forcing. The Vatnajökull ice cap (the second largest in the Northern Hemisphere) is reconstituting itself during periods of climate degradation, at the end of interglacials, under the impact of orbital forcing, but is disappearing during climatic thermal optima. The mild Irminger sea current surrounds the entire island during warm episodes, but drift back to the South during the extension of the East Greenland drift, causing in particular the rapid re-extension of the sea ice and the Vatnajökull ice cap. This consequence makes it possible to track, particularly in Northwest Iceland, the arrival of mild waters from the southern hemisphere, whether they are produced by orbital forcing or abnormal solar activity: these are the Dansgaard–Oeschger events during the Ice Age (with ice shelves) and Bond events during the interglacial period (without ice shelves). A warm Southern Ocean, as at present (second thermal optimum of the Holocene, linked to the optimum of the southern hemisphere), melts the ice caps under the effect of precipitation and contributes to the intensification of volcanism. C.3. The glacial evolution of a volcanic island Iceland today is characterized by landscapes and a sedimentary cover essentially volcanoclastic, inherited from the episodes of glaciations and deglaciations. The largest volumes of lava were emplaced as thick basaltic piles during the Middle Miocene and Pliocene, then eroded by glaciers after a partial gravitational collapse of the Pliocene Icelandic crust. This is corroborated by the significant subsidence of the Northwestern peninsula and the Northeastern coast, flooding the upstream fjords. The Quaternary vegetation cover was thin and regularly disturbed by volcanic episodes. The erosion caused by the glacial episodes is much greater than the volcanic contributions, which remained very localized, especially at the level of the rifts. Isotopic or stratigraphic datings of Quaternary volcanic deposits attest to their limited apparent volume, but still compatible with their production by an active hotspot. These volcanic

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inputs were more cyclic than the Neogene ones, especially during the deglaciation period. The same is true for subglacial volcanoes and hyaloclastite ridges that correlate more or less well with temporary deglaciations (≈ 10 to 20 ky) of very extensive ice sheets on the surface and over time (≈ 100 ky). The propagation and excavation of canyons offshore and onshore at the edges of the oceanic shelf and of the central Icelandic plateau attest to the relative uplift of the latter by mechanical ablation coupled with an isostatic discharge likely to influence the activity of the hotspot since 7 My. The first glaciations, from 9 My until ca 400 ky, were mostly temperate-based and therefore very erosive. It cannot be excluded that the cooling of the upper crust, induced by glaciations and colder oceanic waters since the beginning of the Quaternary, accentuated the thermal subsidence of the margins of the Icelandic Plateau. Unlike Greenland and Scandinavia, the Neogene and Quaternary sedimentary deposits remain relatively thin around the island, on the flooded plateau today at a depth of less than 300 m, shaped by Oligo-Miocene and Pliocene marine abrasion surfaces and regularly cleared by the extensive glaciations. All these characteristics give Iceland and its surrounding environment a unique sentinel status – tectonic, volcanic and climatic – of the northern hemisphere.

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Van Vliet-Lanoë, B., Schneider, J.L., Guðmundsson, Á., Guillou, H., Nomade, S., Chazot, G., Liorziou, C., Guégan, S. (2018). Eemian estuarine record forced by glacio-isostasy (S Iceland): Link with Greenland and deep sea records. Can. J. Earth Sci., 55(2), 154–171. Van Vliet-Lanoë, B., Bergerat, F., Allemand, P., Innocent, C., Guillou, H., Cavailhes, T., Liorzou, C., Grandjean, P., Passot, S. (2020a). Tectonism and volcanism enhanced by deglaciation events in southern Iceland [Online]. Quat. Res., 94, 94–120. Available at: http://doi.org/10.1017/QUA.2019.68. Van Vliet-Lanoë, B., Knudsen, O., Guðmundsson, Á., Guillou, H., Chazot, G., Langlade, J., Liorzou, C., Nonnotte, N. (2020b). Volcanoes and climate: The triggering of Preboreal jökulhlaups in Iceland. Int. J. Earth Sci., 109, 847–876 [Online]. Available at: http://doi.org/10.1007/s00531-020-01833-9. Villemin, T. and Bergerat, F. (2013). From surface fault traces to a fault growth model: The Vogar fissure swarm of the Reykjanes Peninsula, Southwest Iceland. J. Struct. Geol., 51, 38–51. Voight, B., Clifton, A., Hjartarson, A., Steingrímsson, B., Brandsdóttir, B., Rodríguez, C., McGarvie, D., Sigmundsson, F., Ívarsson, G., Friðleifsson, G.O., Larsen, G., Jónsdóttir, G.S., Noll, H., McDougall, I., Kaldal, I., Friðleifsson, I.B., Aronson, J.L., Karson, J.A., Grönvold, K., Young, K.D., Kristjánsson, L., Sigurgeirsson, M.Á., Guðmundsson, M.T., Jancin, M., Flóvenz, Ó.G., Einarsson, P., Williams Jr., R.S., Buck, R., Pálmadóttir, S., Friedrich, W. (2020). A half-century of geologic and geothermic investigations in Iceland: The legacy of Kristján Sæmundsson. J. Volcan. Geoth. Res., 391, 106434 [Online]. Available at: http://doi.org/10.1016/j.jvolgeores.2018.08.012. Werner, R. and Schminke, H.-U. (1999). Englacial vs lacustrine origin of volcanic table mountains: Evidence from Iceland. Bull. Volcan., 60, 335–354. Wolfe., C.J., Bjarnason, I.T., Van Decar., J.C., Solomon, S.C. (1997). Seismic structure of the Iceland mantle plume. Nature, 245–247. Zachos, J., Pagani, M., Sloan, L., Thomas, E., Billups, E. (2001). Trends, rhythms, and aberrations in global climate 65 Ma to present. Science, 292, 686–693. Ziegler, P.A. (1992). North Sea rift system. Tectonophysics, 208, 55–75 [Online]. Available at: http://doi.org/10.1016/0040-1951(92)90336-5. Websites Geological maps, https://en.isor.is.

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(Iceland

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Hydrography (Haffránsoknastoffnun, Marine and Freshwater Research Institute): https://sjora.hafro.is/. Maps and aerial photographs of Iceland (National Land Survey of Iceland; Langmælingar Íslands): https://www.lmi.is/en/. Maps of Iceland: http://map.is. Meteorology, hydrology, seismology, volcanic activity (Icelandic Meteorological Office; Veðurstofa Íslands): https://en.vedur.is/. Vatnajökull National Park: https://www.vatnajokulsthjodgardur.is/en.

List of Authors Françoise BERGERAT

René MAURY

ISTeP CNRS-Sorbonne Université Paris France

Géosciences Océan IUEM Université de Bretagne Occidentale Brest France

Laurent GEOFFROY Géosciences Océan IUEM Université de Bretagne Occidentale Brest France

Hervé GUILLOU

Brigitte VAN VLIET-LANOË Géosciences Océan IUEM Université de Bretagne Occidentale Brest France

Laboratoire des Sciences du Climat et de l’Environnement LSCE/IPSL, CEA-CNRS-UVSQ Université Paris-Saclay Gif-sur-Yvette France

Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

Index A accretion, 177 aerosols, 35, 36, 69, 70 sulfuric, 35, 36, 69 affinity, 77, 80, 170 albedo, 39, 70, 158, 164 Alleröd, 47 allophane, 50, 176 alteration hydrothermal, 72 meteoric, 72 andosols, 175–177 aquifer, 26, 38, 118–120, 123 ash, 26, 27, 29, 32–36, 39, 42, 45, 50, 69, 70, 94, 158

B Barents Sea, 103, 112, 162, 166 basalt, 5, 49, 73, 81, 124, 130 basicrustal, 3, 5 batholite, 7 Bølling, 15, 66, 123, 136, 150, 151, 154, 166, 175 bombs, 31

boreal, 137, 159, 169 break-up, 145 Bronze Age, 170 browning, 176 buoyancy, 164

C caldera, 14, 16, 18, 19, 37, 39, 57, 140, 142, 150 cauldron, 29, 37 chronometer, 71, 72, 80 circulation atmospheric, 33, 70 thermohaline (THC), 109, 114–116, 134, 141, 158 colonization plant, 169, 170 Viking, 18 component, 10–12 cones, 18, 20, 26, 28, 31, 32, 40 rootless, 31 scoriaceous, 18 connection, 17 contamination, 3, 14, 123

Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

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cooling, 25, 26, 31, 34, 35, 47, 48, 73, 80, 94, 106, 109, 110, 115–117, 127, 140, 142, 152, 155, 156, 158, 161, 167, 170 cosmic bombardment, 70 crater, 14, 15, 18, 31, 37 crust, 2, 3, 7, 11–14, 22, 23, 115, 119, 120, 122, 173 continental, 3, 13, 119, 120 cryptogamic, 173 oceanic, 7, 11, 12, 14, 115, 120 crustal, 16 crystals, 70, 72, 76, 78–80, 90, 91, 165 current Irminger, 66, 116, 117, 142, 145, 150, 154, 155, 158, 161, 162, 164, 169, 170, 173 Norwegian, 162 sea, 120

D dacite, 7 dating 40 Ar/39Ar, 71, 75–81, 88–92, 94, 95 tephrostratigraphy, 45, 124, 139 decompression, 42, 67, 68 deep waters, 108, 134, 155 North Atlantic (NADW), 108 degassing, 31, 76, 78 deglaciation, 1, 10, 14–16, 19, 22, 29, 32, 33, 42, 48, 59, 62–70, 86, 88, 91, 94, 116, 120, 123, 125, 130, 133, 136–139, 141–147, 150, 151, 153, 154, 157, 159, 160, 162, 175 degradation, 49, 175

depleted MORB mantle (DMM), 10, 11, 14 diagram, 8, 10, 78 drift East Greeland, 155, 161, 164 North Atlantic, 66, 112, 116, 119, 145, 152, 157, 158, 162, 173 dropstones, 31, 129, 144 ductile, 3 dunes, 50, 51, 178 dwarfism, 170 dyke, 18, 26, 64

E earthquake, 139 Eemian, 42, 66, 137–143, 154, 158, 166, 168, effusive, 17, 60, 66, 68, 80, 123, 126, 130 en échelon, 17, 18 endemic, 168 enrichment, 8 Eocene, 103, 104, 108, 115, 127 Esker, 39 eustatism, 69 event Bond, 42, 119, 158 Dansgaard–Oeschger (DO), 42, 69, 119, 158 Heinrich, 145 exhumation, 106 extensive, 3

F faults, 3, 21, 31, 58, 61, 68 fauna, 34, 35, 167 fissure, 17, 32 float, 37

Index

flood glacial (surge), 39, 119, 148, 149 flora, 127, 167, 169, 170 flow, 37, 40, 42, 44, 49, 50, 59, 117, 130, 148 laminar, 50 turbulent, 42 forcing, 69, 137, 150, 158, 164, 170 orbital, 69, 137, 150, 158, 164, 170 forebulge, 122 marginal, 122 fractures, 4, 6, 39, 57 fracturing, 68 fragile, 3 frasil, 165 fusibility, 37

G gas, 18, 26, 28, 33, 35, 36, 69, 72, 74–76, 78 Gelasian, 159 Geothermal, 7, 31, 38, 176, GIA (glacio-isostatic accommodation), 122 GIFR (Greenland–Iceland–Faroe Ridge), 104 GIR (Greenland–Iceland Ridge), 115 glacial advance, 122, 139, 140, 143, 154 glaciation, 110, 114, 123, 133, 137, 142, 143, 146 glacier surging, 118, 119 graben, 29 gravitational spreading, 66, 118, 119 growth, 14, 32, 57, 72, 75, 110, 123 Gulf Stream, 110, 161, 162, 164, 170

217

H halloysite, 176 Holocene, 3, 18, 23, 34, 46, 49, 66, 68, 69, 110, 114, 119, 120, 124, 140, 150, 152, 154, 158, 170, 176 hot spot, 1, 3, 5, 13, 15, 20, 33, 66, 112, 115, 116, 126, 161 hyaloclastites, 64, 65, 159 laminated, 64, 65 hydrated glass, 26, 31 hydrolaccolith (pingo), 174 hydromagmatic, 1, 28 hydrostatic pressure, 39, 73, 117

I ice cap, 47, 58, 70, 114, 127, 147, 148 cores, 33, 44, 45, 47, 70, 89, 94, 105, 108, 126, 137, 139, 140, 142, 143, 145 sheet, 65, 110, 112 stream, 112, 117, 148 Ice Age, 15, 42, 61, 63, 69, 94, 114, 158 Little, 36, 42, 66, 119, 155, 157, 158, 166, 167, 175, 178 iceberg, 39, 153 Icelandic cyclonic depression, 161 ICPMS (inductively coupled plasma mass spectrometry), 47 igneous province, 13 ignimbrite, 89, 90, 94, 145 imogolite, 50, 176 inflation, 18, 49 insularization, 115 intercept, 78, 90 interstadial, 94, 137 intrusion, 83, 164

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IRD (ice-rafted debris), 105, 108, 110, 119, 133, 134, 136, 144, 145 irradiation, 76, 78, 80 isochron, 77–80, 90, 91 isostatic rebound, 19, 120, 155

Miocene, 104, 106, 108, 115, 127, 159, 169 Móbergs, 57, 58, 73, 81 MORB (mid-ocean ridge basalts), 3, 4, 8–10

J, K, L

N

jökulhlaup, 18, 35, 37–44, 46, 47, 49–51, 119, 120, 137, 139, 158, 175, 178 kaolinite, 176 lapilli, 31, 45 Last Glacial Maximum (LGM), 66, 125, 135, 137, 144–147, 149, 166 lava aa, 21, 23, 24, 81 flow, 24, 31, 172 lake, 20, 23, 60 pahoehoe, 21, 22 pillow, 26 ropy, 21–23, 25 liquefaction, 176 lithalsa, 172–175 lithosphere, 3, 67, 69, 120 Little Ice Age, 36, 42, 66, 119, 155, 157, 158, 166, 167, 175, 178

NADW (North Atlantic Deep Water), 108 NAO (North Atlantic oscillation), 119, 161, 167 native, 170 Neogene, 109, 114, 115, 169 Neoglacial, 66, 156 non-Newtonian liquid, 40 nunataks, 146, 170

M magma, 1, 3, 5, 13, 14, 18, 21, 22, 26, 28, 45, 60, 62, 64–66, 69, 72, 73, 77, 80, 88, 125, 161 magnetic field, 74 mantle, 3, 8, 10–14, 57, 67, 69, 120 melting, 67 partial, 7, 10, 57, 65, 67, 69 megadune, 42 Messinian, 196 Middle Pleistocene Transition (MPT), 112, 130

O obsidian, 90 ocean Arctic, 106, 110, 112, 145, 158 Southern, 114, 158 OIB (ocean island basalts), 6, 8–11 Oligocene, 103, 106, 115, 169 optimum climatic, 168, 170 Medieval, 157 thermal, 116, 139, 154, 155, 165, 167, 175

P pack ice, 70, 106, 109, 110, 112, 114, 116, 117, 136, 140, 143, 145, 150–152, 155, 157–159, 161, 164–167 Papar, 168 passive margins, 115 peat, 168 bog, 175

Index

periglacial, 126 permafrost, 29, 117, 142, 155, 157, 173–175, 177 phreatomagmatic, 16, 18, 19, 29, 30, 33 pioneer, 170 Pleistocene, 109, 110, 112, 123, 126, 128, 130, 135, 144, 169 Plinian, 16, 33, 139, Pliocene, 109, 112, 114, 115, 126, 169 plume, 3, 11–14, 16, 27, 33 progradation, 59 propagation, 22 prostration, 170 pumice, 34, 35

Q, R Quaternary, 114, 123, 133, 134 radioactive decay, 80 radiocarbon (carbon 14), 71, 123, 124 radiometric ages, 86, 87 regression, 123, 139, 169 reservoir, 5, 16, 18, 20, 58, 71, 123, 125 magma, 5, 18 residence time, 88 rhyolite, 76, 90, 92, 94, 95 ridge, 3, 4, 115 hyaloclastite, 64, 65, 134–136, 159 Mid-Atlantic (MAR), 3, 126 oceanic, 31, 116 rift jump, 109, 110, 127, 128 rifting, 13, 20, 104, 106, 108, 115 rocky, 143, 144, 150, 157, 173, 175 rofaborð, 50, 178

219

S sandur, 18, 39, 40, 43, 150 Scandinavia, 110, 115, 116, 120, 135, 155, 167 SDRs (seaward-dipping reflectors), 196 sea level, 107, 110, 119, 120, 122, 135, 139, 144, 145, 148–150, 154 seasonal frost, 174–177 sediment, 31 seismic activity, 18 crisis, 18, 135 seismogenic, 2 Serravallian, 115, 127 shear, 39, 123 shelf, 104, 108, 110, 112, 118, 119, 123, 130, 137, 142–145, 148, 150, 159, 160, 165 sill, 64 snowy, 170 solar cycle, 158 minimum, 175 solifluction, 48, 177 stadial, 94, 136, 137 steam, 23, 26, 27, 31 storminess, 170 strandflat, 130, 139, 152, 159 stranding, 170 stratigraphy, 33, 94, 95, 123, 126, 128 sequential, 123 stratovolcano, 14, 21 stress, 14, 58, 68 crustal, 68 sub-glacial, 1, 16, 18, 26, 28, 38–40, 42, 57–61, 70, 81, 83, 84, 86, 87, 90, 92, 117, 120, 125, 126, 130, 135, 137, 143, 144, 146, 159

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subsidence, 106, 109, 115, 120, 143, 159, 169 supercritical, 40 suture, 2 swarms, 3, 58, 61

T tectonics, 6, 16, 108, 116, 123, 169 terminal moraine, 39 termination, 22, 152 terrestrial obliquity, 69, 110 thermal contraction, 175 thermokarst, 175 thixotropic, 176 tindar, 57, 61 Tortonian, 107, 169 transform, 31 transgression, 29, 122, 123, 140, 142, 144, 153 transtensive, 115 tumuli, 22 tundra, 127, 170, 172, 178 tuyas, 57, 58, 61, 62, 70, 81, 90, 125, 126, 130, 134, 135, 140, 144, 151, 159

U, V unloading, 1, 20, 39, 65–68, 134, 136, 161 varves, 70, 123 vesicles gas, 26 viking, 18, 168 volcano central, 18, 142 fissural, 31 shield, 14, 15, 150

Z zone Tjörnes Fracture (TFZ), 4, 31 transform, 31 volcanic, 16, 66, 69, 89, 142, 150 east (EVZ), 4, 16, 19 north (NVZ), 4, 31, 159 west (WVZ), 89

Summary of Volume 1 List of Abbreviations Preface Brigitte VAN VLIET-LANOË and Françoise BERGERAT Introduction Brigitte VAN VLIET-LANOË and René MAURY Chapter 1. Iceland, in the Lineage of Two Oceans Brigitte VAN VLIET-LANOË and Françoise BERGERAT 1.1. Geographic and geodynamic context 1.2. Components of the North Atlantic domain 1.2.1. The Mid-Atlantic Ridge 1.2.2. The North Atlantic Igneous Province 1.2.3. The Icelandic hot spot 1.2.4. The Greenland–Iceland–Faroe Ridge 1.3. Geodynamic characteristics of Iceland 1.3.1. Seismicity 1.3.2. Icelandic volcanism 1.3.3. Eustatism and the Icelandic glaciers 1.4. References

Iceland Within the Northern Atlantic 2: Interactions between Volcanoes and Glaciers, First Edition. Brigitte Van Vliet-Lanoë. © ISTE Ltd 2021. Published by ISTE Ltd and John Wiley & Sons, Inc.

Iceland Within the Northern Atlantic

Chapter 2. Iceland, an Emerging Ocean Rift Françoise BERGERAT 2.1. Mid-Atlantic Ridge and Icelandic hot spot interactions 2.2. Present-day deformations in Iceland 2.2.1. Seismicity 2.2.2. Motions at plate boundaries 2.3. Iceland’s main structural features 2.3.1. The paleo-rifts and the active rift 2.3.2. The transform zones 2.4. Geothermal energy and hydrothermalism 2.4.1. Geothermal systems 2.4.2. Geysers and hydrothermalism 2.5. References Chapter 3. Iceland, A legacy of North Atlantic History Laurent GEOFFROY 3.1. Bathymetry of the Northeast Atlantic domain and geoid anomalies 3.2. The North Atlantic and the continental breakup of Laurussia 3.2.1. Passive margins and large igneous provinces 3.2.2. The early beginnings of the opening of the North Atlantic Ocean 3.2.3. Thulean magmatism in the Paleocene and the continental breakup of the Northeast Atlantic 3.2.4. Chronology and kinematics of the opening of the Northeast Atlantic 3.2.5. The Northeast Atlantic region: mantle plume or not? 3.3. The origin of Iceland 3.3.1. The anomalous crust of the GIFR ridge and the deep structure of Iceland 3.3.2. Icelandic SDRs 3.3.3. Interpretations of GIFR and Iceland 3.4. References

Summary of Volume 1

Conclusion Françoise BERGERAT and Laurent GEOFFROY References