Faculty of Science and Technology Department of Geosciences Holocene glacier history of Svalbard: Retracing the style of (de-)glaciation — Wesley R. Farnsworth A dissertation for the degree of Philosophiae Doctor – December 2018 Holocene glacier history of Svalbard: Retracing the style of (de-)glaciation Wesley R. Farnsworth iii Holocene glacier history of Svalbard: Retracing the style of (de-)glaciation Wesley Randall Farnsworth A dissertation submitted to the Faculty of Science and Technology, UiT The Arctic University of Norway for the degree of Philosophiae Doctor (PhD) Department of Arctic Geology The University Centre in Svalbard (UNIS) And Department of Geosciences UiT, The Arctic University of Norway December 2018 iv © Wesley Randall Farnsworth, 2018 Title: Holocene glacier history of Svalbard: Retracing the style of (de-)glaciation Author: Wesley Randall Farnsworth Supervision by: Prof. Ólafur Ingólfsson (University Centre in Svalbard & University of Iceland) Prof. Michael Retelle (Bates College & University Centre in Svalbard) Prof. Anders Schomacker (UiT, The Arctic University of Norway) All rights reserved. No part of this publication may be reproduced or transmitted, in any form or by any means, without permission. ISBN (print): ISBN (online): Print in Norway: Front cover illustration: Modified oblique aerial photograph (ID: S38_1850) taken in 1938 by the Norwegian Polar Institute from an altitude of 2900 m a.s.l. above the eastern margin of Valhallfonna with a view to the south down Lomfjorden. The bay of De Geerbukta and the Fakse moraine system are located in the foreground. v Abstract Identifying the key factors that influence the global cryosphere and eustatic sea level are critical in today’s populated world characterized by a changing climate. Within the Arctic, which has recently experienced amplified warming, and located between the Polar Northern Atlantic and the Arctic seasonal sea-ice, the Svalbard archipelago experiences a heightened sensitivity to climate change. Studying the processes, dynamics and historic fluctuations of Svalbard’s glaciers and climate is critical. Understanding these elements allows us to place current rates of change into longer-term perspective and ultimately to better model future climatic conditions. This study synthesizes the state of the art of Svalbard’s Holocene glacial and climate history. Chapters (i) introduce new findings of Holocene glaciers and climate; (ii) discuss the factors influencing Svalbard ice margins; (iii) summarize accumulated knowledge in the perspective of today’s paradigm; and (iv) outline potential approaches to address further unknowns regarding the Holocene on Svalbard. Through the Holocene, Svalbard glaciers have exhibited at least two phases of widespread re-advances, one during the Early Holocene and another throughout the entire Late Holocene. No geomorphological features have been identified corresponding to glacier re-advances between 9.0 – 4.5 ka BP. The Early Holocene glacier re-advances are identified across Svalbard and correspond to a diverse range of glacier sizes. With our current level of age constraint, these ice marginal fluctuations do not appear synchronous. Furthermore, the Early Holocene climate is believed to have been warm, unfavorable for glacier growth, and characterized by deglaciation. Early Holocene glacier re-advances appear to relate to the time-transgressive nature of deglaciation. Thus, the re- advances correspond to glacio-dynamics (not mass balance) and reflect the complex style of ice- mass-loss during a changing climate. Landforms and deposits from glaciers re-advancing during the Late Holocene have been the primary focus of Holocene glacial studies. Glacier re-advances and corresponding deposits have been attributed to episodic Neoglacial cooling and the Little Ice Age (LIA). The majority of Late Holocene glacier re-advances have been dated to between 4.0 – 0.5 ka BP with the highest frequency of re-advances constrained to 1.0 – 0.5 ka BP, during the first half of the LIA. It has been suggested that glacial landforms and deposits from LIA re-advances indicate rapid and dynamic glacier behavior, and in some cases surge-type events. During the 20th century (i.e. post-LIA), Svalbard glaciers have exhibited widespread negative mass balance, ice marginal retreat, and glacier thinning. This phase of retreat has had a direct influence on glacier thermal regime, hydrologic system and surface profile. Through the 20th century, some Svalbard glaciers have continued to exhibit surge-type re-advances. Several glaciers have surged numerous times. These glacio-dynamic re-advances have been un-sustained and each subsequent surge has been less extensive then prior surges. Consequently, and despite re-advance, glaciers reflect a continual phase of ice-mass-loss in a periodic fashion. Our understanding of Svalbard’s Holocene glaciers and climate has progressed but critical components remain obscure. For example, although our understanding of the timing of the Holocene glacial minimum has improved, we lack detailed constraints on the extent of ice retreat across Svalbard during the Mid-Holocene. As reconstructions of palaeo-temperatures develop, the improvement of palaeo-precipitation proxies (e.g., leaf wax hydrogen isotopes) should continue. Additionally, as we approach further unknowns of Svalbard’s Holocene history, it is evident that studies must take a holistic approach. Combining a mixture of archives, geochronological methods and emerging techniques will enhance the accuracy of reconstructions detailing Svalbard’s glacial history. vi Acknowledgements Working under the supervision of Ólafur Ingólfsson, Michael Retelle and Anders Schomacker has been a highlight of this project. The experience, knowledge and encouragement I have been exposed to through last years in this team has exceeded all belief. I feel so fortunate to have the support, guidance and perspective of such an inspiring group of scientists, educators and leaders. During my doctoral, I have also had privilege to be surrounded by outstanding mentors and role models. I am grateful for the boundless motivation, support and discussions fostered by; Helena Alexanderson, Ívar Örn Benediktsson, Skafti Brynjólfsson, Anders Bjørk, Max Eckerstorfer, Mark Furze, Graham L. Gilbert, Sten-Andreas Grundvåg, Anne Hormes, Ole Humlum, Kristian Kjeldsen, Kurt Kjær, Harold Lovell, Marc Macias-Fauria, Jordan Mertes, Hreggviður Norðdahl, Lena Rubensdotter, Heïdi Sevestre and Alan Werner. Thank you. The development of this doctoral thesis has been greatly aided by a group of individuals whom have donated time, data, energy and expertise. I would like to thank all co-authors for their contributions and improvements to the manuscripts. Additionally, the thesis text has benefited from constructive feedback from Graham L. Gilbert, David C. Farnsworth and supervisors. I have been fortunate enough to develop this project in the Arctic Geology Department at the University Centre in Svalbard. The hallway is my home and I am grateful for the community that evolves within it. Thank you for the company, encouragement, coffees (beers) and cakes. This doctoral project has also benefited from great support in the field (Sveinn Brynjólfssson & Sara Mollie Cohen), and internally at UNIS; from the Library (Berit Jakobsen & Catherine Fjeldstad) and from Logistics. Thank you for your support. This project has also greatly benefited from the open doors and inviting research communities at The Natural History Museum of Denmark, University of Copenhagen and the School of Engineering and Natural Science at the University of Iceland. This doctoral project has been generously supported by the UNIS internal research fund, Svalbard Environmental Protection Fund, Carlsberg Foundation, Arctic Research and Studies, ResClim Research School, Arctic Field Grant from the Svalbard Science Forum and Letterstedtska Föreningen. The ever-evolving community of Svalbard friends that has anchored me since January 2010, up until today, is the most valuable aspect of my life up here. I cannot image these last years without the love, laughs and strength that has come from you all. We have also had a pretty nice time. Thank you. Finally, I would like to thank my family for their limitless love, support and encouragement. I also am endlessly grateful for my partner, Hrefna Dögg Gunnarsdóttir, for her love, patience and balance. Thanks to you all. vii Preface This dissertation is submitted in partial fulfillment of the requirements for the Degree of Philosophiae Doctor (PhD) in Science. This thesis is a product of a four-year PhD study carried out at the University Centre in Svalbard (UNIS) between December 2014 and December 2018. The doctoral project was conducted in collaboration with UiT The Arctic University of Norway in Tromsø. The project was supervised by Ólafur Ingólfsson (University of Iceland and UNIS), Michael Retelle (Bates College and UNIS) and Anders Schomacker (UiT and the Natural History Museum of Denmark, University of Copenhagen). While the four-year doctoral position was funded by the University Centre in Svalbard, additional funding for field work, data analysis, conference attendance and mobility was sought through different funding agencies. This doctoral project was financially supported primarily by the UNIS internal research fund (to Ingólfsson), Svalbard Environmental Protection Fund (16/35 to Farnsworth), Carlsberg Foundation (CF14-0756 to Schomacker), Arctic Research and Studies (to Schomacker & Farnsworth), ResClim Research School (to Farnsworth), Arctic Field Grant from the Svalbard Science Forum (to Schomacker) and Letterstedtska Föreningen (to Farnsworth). In fulfillment of the 25% teaching and duty work written into the doctoral contract, Farnsworth gave lectures, ran exercises, instructed during field excursions and assisted in eight UNIS courses. Since spring 2015, Farnsworth contributed to 13 different classes including bachelor courses AG-204, AG-210, AG-220, as well as graduate courses AG-326/826, AG-330/830, AG- 346, AG-348/848 and AS-301. In addition to teaching, research findings from this doctoral project were disseminated through oral and poster presentations at 12 international conferences throughout the Nordic countries and North America. Oral presentations were given at all conferences unless otherwise noted. 2018: PAST Gateways Durham, United Kingdom April; 33rd Nordic Geological Winter Meeting Copenhagen, Denmark January; 2017: Geological Society of America, GSA Annual Meeting Seattle, USA October; International Quaternary Webinar October; PAST Gateways Kristineberg, Sweden May; 2016: PAST Gateways Trondheim, Norway May; 32nd Nordic Geological Winter Meeting Helsinki, Finland January; 2015: American Geophysical Union Annual Meeting, AGU San Francisco, USA December (Poster); ResClim PhD. Forum Askö, Sweden November; Nordic International Glaciological Symposia, IGS Copenhagen, Denmark October; International Glaciological Society, IGS Symposia Höfn, Iceland June (Poster); 45th International Arctic Workshop Bergen, Norway May (Poster). viii This doctoral thesis is a synopsis of five research manuscripts: *Farnsworth, W.R., Ingólfsson, Ó., Retelle, M., Allaart, L., Håkansson, L., Schomacker, A. (2018): Svalbard glaciers re-advanced during the Pleistocene-Holocene transition. Boreas 47, 1022-1032. DOI:10.1111/bor.12326. **Larsen, E., Lyså, A., Rubensdotter, L., Farnsworth, W.R., Jensen, M., Nadeau, M. Ottesen, D. (2018): Lateglacial and Holocene glacier activity in the Van Mijenfjorden area, western Svalbard. arktos. DOI: 10.1007/s41063-018-0042-2. *Farnsworth, W.R., Ingólfsson, Ó., Noormets, R., Allaart, L., Alexanderson, H., Henriksen, M., Schomacker, A. (2017): Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard. Boreas 46, 585-603. DOI: 10.1111/bor.12269. ***Farnsworth, W.R., Ingólfsson, Ó., Schomacker, A., Retelle, M. (2016): Over 400 previously undocumented Svalbard surge-type glaciers identified. Geomorphology 264, 52-60. DOI: 10.1016/j.geomorph.2016.03.025. Farnsworth, W.R., Ingólfsson, Ó., Schomacker, A., Retelle, M. in prep. Holocene glacial and climate history of Svalbard - status, perspectives and challenges. Reprints were made with permission from *John Wiley & Sons, **Springer and ***Elsevier. ix CONTENTS 1. INTRODUCTION 1 1.0.1 The Holocene 2 1.1 Svalbard 2 1.1.1 Regional setting and modern climate 2 1.1.2 Glacial history and relative sea level 3 1.1.3 Landscape and glaciers 6 1.1.4 Surge-type glaciers, behavior and landform assemblages 8 1.1.5 Aims and research objectives 9 2. METHODS 11 2.0.1 Terrestrial stratigraphy and mapping 11 2.0.2 Marine data acquisition and processing 11 2.0.3 Sampling and radiocarbon dating 12 3. AUTHOR CONTRIBUTIONS AND MANUSCRIPT SUMMARIES 13 3.1 Chapter I Farnsworth et al. 2018 14 3.2 Chapter II Larsen et al. 2018 15 3.3 Chapter III Farnsworth et al. 2017 16 3.4 Chapter IV Farnsworth et al. 2016 17 3.5 Chapter V Farnsworth et al. in prep 18 4. DISCUSSION 19 4.1 Svalbard Holocene climate and the unknowns 19 4.1.1 Atlantic waters control the temperature 19 4.1.2 Palaeo-precipitation remains unknown 19 4.2 Drivers of Holocene glacier re-advances 20 4.2.1 Early Holocene re-advances 20 4.2.2 Neoglacial – Little Ice Age re-advances 21 4.2.3 Post-Little Ice Age re-advances 22 4.2.4 Early vs. Late Holocene ice cover variations 23 4.3 Advances in deglaciation 24 4.3.1 Restrained rebound and ice caps chasing the equilibrium line 24 4.3.2 Late Pleistocene – Early Holocene ice shelves on Svalbard 27 4.3.3 Modern glacier behavior as a model for past glacier dynamics 28 5. SUMMARY AND CONCLUSIONS 33 6. FUTURE PERSPECTIVES 35 References 36 Research papers Chapter I Farnsworth et al. 2018 Chapter II Larsen et al. 2018 Chapter III Farnsworth et al. 2017 Chapter IV Farnsworth et al. 2016 Chapter V Farnsworth et al. in prep Appendices 187 x 1 1. INTRODUCTION Present geologic processes provide insight into past geological events, and a window into the future (Lyell 1830; Hume 1955). Therefore, studying glaciers and their history is crucial for extending our relatively short observational record, putting our process understanding into a long-term perspective, and better anticipating future climatic conditions (Ingólfsson & Landvik 2013; Hughes et al. 2016). Arctic glaciers are valuable indicators of past climate variations given their sensitivity to winter precipitation and summer temperatures (Oerlemans 2005). Investigations of past changes in high-latitude glaciers and climate allow us to better understand the role of the Arctic in the global climate system (McKay & Kaufman 2014). Today, climatic shifts are occurring with an amplified effect in the Arctic and did so following the last deglaciation as well (Hald et al. 2007; Hartmann et al. 2013). Current climatic conditions are driving negative glacier net mass balance and reduction in the global cryosphere (Kaser et al. 2006; Kjeldsen et al. 2015; Huss & Hock 2015). These changes in ice volume have a direct impact on global sea level (DeConto & Pollard 2016; Shepherd et al. 2018; Bamber et al. 2018). Understanding the dynamics that govern the stability of marine based ice sheets like the West Antarctic Ice Sheet (WAIS) is critical, given the (potential) sea level equivalent stored within them (Schoof 2007; Joughin et al. 2014). While we now have decadal-scale observational records for modern ice sheets, the past ice sheet records can offer a detailed long-term perspective on the style and characteristics of ice sheet behavior (Patton et al. 2015; Gandy et al. 2018). The Svalbard Barents Sea Ice Sheet (SBSIS) is a palaeo-analogue for the WAIS and can provide valuable insights into the dynamics and style of deglaciation that can be expected from a marine-based ice sheet in a rapidly changing climate (Mercer 1969, 1970; Patton et al. 2015, 2017; Esteves et al. 2017). Understanding the processes, controls, and dynamics acting on the SBSIS during the Late Pleistocene and Early Holocene allow us to better predict the future evolution of the WAIS (Winsborrow et al. 2012). Investigations of ice sheet dynamics commonly target ice streams, the rapidly flowing rivers of ice that drain disproportionately large volumes of the ice sheet (Stokes & Clark 1999; Bennett et al. 2003; Briner 2016; Stokes et al. 2016; Larsen et al. 2018). Ice stream behavior and dynamics have been investigated in contemporary glacier systems and reconstructed from palaeo-drainage networks based on their geomorphological traces (Andreassen & Winsborrow 2009; Winsborrow et al. 2012; Joughin et al. 2014; Kleman & Applegate 2014; Margold et al. 2015; Stokes et al. 2016). Distinct landsystem models have been developed to identify and categorize past ice streams (Stokes & Clark 1999; Evans et al. 2008; Andreassen et al. 2014). Interestingly, similarities exist between 2 the geomorphological assemblages developed for ice streams and surge-type glaciers, which may indicate a dynamic commonality (Kjær et al. 2008; Winsborrow et al. 2012; Andreassen et al. 2014; Ingólfsson et al. 2016; Newton & Huuse 2017). Svalbard and the surrounding marine environment are an ideal region to investigate glaciers and climate history. The terrestrial and marine archives are fingerprinted with the final stages of deglaciation of the marine based SBSIS (Ingólfsson & Landvik 2013, 2014; Flink et al. 2017; Hogan et al. 2017). Furthermore, investigations of Svalbard glaciers are particularly important due to the region’s climatic sensitivity, amplified response to global shifts in atmospheric and oceanic temperatures, as well as the high density of documented surge-type glaciers (Jiskoot et al. 2000; Benestad et al., 2003; Hagen et al. 2003; Hald et al. 2007; Sevestre & Benn 2015; Isaksen et al. 2016). 1.0.1 The Holocene The end of the Pleistocene, characterized by a cold period known as the Younger Dryas (YD), ended at 11.7 ka BP and transitioned rapidly into the warmer Holocene period (Dansgaard et al., 1993; Cohen et al. 2013; updated). The global climate during the Holocene has traditionally been regarded as relatively stable, compared to the preceding Late Pleistocene (Dansgaard et al. 1993; Steffensen et al. 2008; Rockström et al. 2009). However, this paradigm of a relatively uneventful Holocene is being increasingly challenged (Bond et al. 2001; Mayewski et al. 2004; Wanner et al. 2011). The Holocene period is separated into three sub-divisions or stages; Early, Mid and Late Holocene, which correspond to the Greenlandian, Northgrippian and Meghalayan stages, respectively (Walker et al. 2012; Cohen et al. 2013; updated). The timing of these stages is partitioned as follows: from 11.7 – 8.2, 8.2 – 4.2 and 4.2 to present. In this text, the Late Holocene stage is, in turn, broken into three time periods, Neoglacial (4.2 ka BP – 1920 AD), Little Ice Age (LIA; 1250 – 1920 AD), and Post-LIA (1920 – present). 1.1 Svalbard 1.1.1 Regional setting and contemporary climate Located along the dominant corridor of atmospheric moisture between the Atlantic and the Arctic Basin, the Svalbard archipelago spans from 74° – 81° N and 10° – 35° E (Fig. 1; Drange et al. 2013). The region has a sensitive climate due to its position at the northern extent of the North Atlantic Drift (West Spitsbergen Current; Fig. 1) and the southern border of the Arctic sea-ice front (Rogers et al. 2005). Svalbard is categorized as having a dry, high-Arctic climate with periglacial conditions, extreme winter temperatures and warm continuous permafrost (French 2007; 3 Christiansen et al. 2010). Despite its northerly latitude, Spitsbergen, the largest of Svalbard’s islands, (followed in size by Nordaustlandet, Edgeøya and Barentsøya), currently experiences a relatively mild climate. The warm West Spitsbergen Current travels off the western extent of Svalbard and influences weather patterns and sea-ice distribution (Fig. 1B; Førland et al. 1997). Regional climate is controlled by the interactions between the Icelandic Low and the Siberian High pressure systems where high temperatures (and precipitation) are driven north over Svalbard by the North Atlantic cyclone track (Hanssen-Bauer et al. 1990; Humlum 2002). Svalbard precipitation is closely coupled to the mode of the North Atlantic Oscillation (Dickson et al. 2000) and falls predominantly in solid form. The interactions of these air masses along the western flank of Svalbard result in relatively warmer and wetter winter conditions than are typical for such latitudes (Førland et al. 1997; Eckerstorfer & Christiansen 2011). Over the last century, most of the annual variability seen in mean annual air temperature (roughly 4 – 5 °C on an annual – decadal scale) is a result of fluctuations occurring during the winter and shoulder-season months (September – November and March – May). Summertime (June, July and August) temperature averages since the start of the Longyearbyen record in 1912 have gradually increased from roughly 4 °C to 5.5 °C, but exhibit minimal variability, i.e. in between 0.5 – 1.0 °C (Christiansen et al. 2013). 1.1.2 Glacial history and relative sea level Today, it is widely accepted that the Svalbard Barents Sea region has undergone repeated glaciations through the Late Quaternary (Mangerud et al. 1998; Landvik et al. 1998; Svendsen et al. 2004; Larsen et al. 2006). The concept of an ice sheet covering Svalbard, the Barents Sea and extending as far south as Scandinavia has developed for over a century (De Geer 1900). In 1860, raised marine beaches with fossil bearing marine sediments were described in Norway and were suggested to relate to uplift of the earth’s crust following the unloading of an ice mass (Kjerulf & Sars 1860). In the late 1800s raised marine beaches with varying maximum elevations were likewise observed across Svalbard, in some locations extending up to 100 m a.s.l. (Pike 1898; Nathorst 1899, 1901; Salvigsen 1981). The early studies that dated these raised marine shorelines and examined the signatures of relative sea level changes form the foundation of current understanding of the former center of ice load over Svalbard and the Barents Sea as well as the rates of glacio-isostatic rebound (Feyling- Hanssen & Olsson 1960; Blake 1961, 1962; Schytt et al. 1968; Hoppe 1972). More recent radiocarbon dated relative sea level curves from across Svalbard and the Barents Sea have enhanced our understanding of the pattern of uplift introduced by the earlier studies (Salvigsen 1981; Forman 1990; Bondevik et al. 1995; Landvik et al. 1998; Forman et al. 2004). 4 Fig. 1 A) Map of North Atlantic with the most-credible time-slice reconstructions (18, 16, 14 & 12 ka BP) of ice cover over Scandinavia, Svalbard and the Barents Sea from DATED1 (modified from Hughes et al. 2016). Extensive ice retreat occurs around Svalbard between 15 – 10 ka BP. B) Bathymetry of the North Atlantic, with Svalbard located at the northwestern extent of the Eurasian plate and the Barents Sea. The warm West Spitsbergen Current (WSC) runs up along Svalbard’s western margin while the cool East Spitsbergen Current (ESC) traces down and around from the east (map modified from IBCAO). C) Topographic map of Svalbard with place names. Site 5 locations presented in Chapters I – III. Chapters IV and V target the whole region. The islands of Kvitøya located to the northeast of Nordaustlandet and Bjørnøya located between Norway and Spitsbergen are not shown on the map (modified from the Norwegian Polar Institute). Although the maximum elevations of raised marine sediments outline a general region influenced by substantial ice cover, large uncertainties remained around the geometry and extent of the ice. For decades, investigations focused on and debated the extent of ice cover across Svalbard during the Late Weichselian or Last Glacial Maximum (25 – 15 ka BP; LGM; Ingólfsson & Landvik 2013; Hughes et al. 2016). Opposing positions developed regarding the recent glacial history of the Svalbard region. The “maximalist” school of thought argued extensive ice cover based on correlations of terrestrial and marine stratigraphic records, suggesting Weichselian ice reached the shelf west of Svalbard with limited ice-free areas existing (Mangerud et al. 1992; Mangerud & Svendsen 1992; Svendsen et al. 1992, 1996). The “minimalists” argued for substantially less extensive LGM ice, based on the apparent lack of Late Weichselian glacial deposits as well as the preservation of sediments and landforms of pre-LGM age which had been identified across the west coast of Spitsbergen (Salvigsen 1977; Boulton 1979; Miller 1982; Forman & Miller 1984; Lehman & Forman 1987; Forman 1989; Houmark-Nielsen & Funder 1999; Andersson et al. 1999, 2000). As additional marine geological data have been collected for the region, it has become evident that the extent of LGM ice cover was closer to the view supported the maximalist reconstructions. However, the observations of preserved old landforms by the minimalists reflects the complexity and dynamics of the SBSIS fingerprinted between its regions of streaming and non- streaming ice (Lehman & Forman 1992; Mangerud et al. 1992; Landvik et al. 1998; Ottesen et al. 2007; Landvik et al. 2005, 2013; Ingólfsson & Landvik 2013). At present, there is a firm understanding of the extent of Late Weichselian glaciation and general timing of deglaciation (Fig. 1A; Hormes et al. 2013; Hughes et al. 2016), but knowledge of the thickness, dynamics and behavior remain less clear (Landvik et al. 2014). While the ice thickness of the SBSIS is unknown, submarine glacial landforms and geomorphological evidence indicate grounded ice identified beneath 500 m water depth suggest minimum ice thickness values (Landvik et al. 1998). Numerical ice sheet models suggest ice thickness on the order of 1500 – 3000 m thick (Lambeck 1995; Patton et al. 2015). As the SBSIS is often compared to the WAIS, measured ice thickness on the contemporary ice sheet can be used as a gauge for the potential ice thickness over the SBSIS during the Late Weichselian (Mercer 1970). In the West Antarctic grounded ice has been identified 2 km beneath contemporary sea level and ice thicknesses range from 1.5 – 3 km thick (Fretwell et al. 2012; Jamieson et al. 2014). As the current vision of the SBSIS has developed, reconstructions of the size, shape and behavior of the ice sheet have evolved. What was once a concentric mono-domed slow-moving ice 6 sheet, regulated by large-scale northern hemisphere climatic oscillations, has become a multi- domed, dynamic ice sheet characterized by regions of streaming ice with intermittent zones of non- erosive slow creeping ice cover. Furthermore, the complex behavior is influenced by not just climate fluctuations, but additional factors including relative sea level, subglacial topography, substrate characteristics, basal temperature, and hydrological conditions (Ingólfsson & Landvik 2013). 1.1.3 Landscape and glaciers Svalbard is 62,000 km2 of glaciated mountainous terrain interrupted by low-lying open fjord-valleys leading into over-deepened fjords systems (Ottesen et al. 2007; Gilbert et al. 2018). Svalbard exhibits geologic strata ranging from Precambrian to Quaternary (Hisdal 1985; Dallman et al. 2015). Lithologic and structural controls govern topography and mountain morphology. Topography of the western and northern coasts of Spitsbergen are characterized by high relief alpine terrain while the central and northeastern regions exhibiting gentle plateau-shaped highlands with intersecting incised valleys (Humlum 2002; Dallman et al. 2015). Large portions of northern Spitsbergen exceed 1200 m a.s.l. with many peaks extending over 1600 m a.s.l., while mountain summits further south range around 1000 m a.s.l. The island of Nordaustlandet reaches up to between 600 – 700 m a.s.l. while peaks on Barentsøya and Edgeøya are closer to 500 m a.s.l. (Fig. 1C). Strand-flats and low lying regions are overlain by post-glacial raised marine sediment are commonly found along the outer coasts of Svalbard (Fig. 2A). Higher elevations are generally ice covered with exceptions existing around interior regions of Spitsbergen, which are located in a precipitation shadow (Humlum 2002). Glaciers cover roughly 57% of Svalbard at present and over 65% of that area drains into tidewater glaciers (Blaszczyk et al. 2009; Nuth et al. 2013). The total volume of Svalbard’s glaciers is estimated at 6700 ± 835 km3 or 17 ± 2 mm of sea level equivalent (Martin-Español et al. 2015). Estimates derived from aerial imagery from the early 1990s suggest ice cover ranged around 60% (Hagen et al. 1993) while estimates of glacier extent during the end of the LIA c. 1920, indicate over 70% of Svalbard was glaciated (Martin-Moreno et al. 2017). Glacier types on Svalbard include cirque, valley, and fjord glaciers systems, thin alpine ice caps and extensive ice caps, unconstrained by regional topography and grounded below sea level (Fig. 2B, C & D). 7 Fig. 2 Photo-mosaic of Svalbard landscape and glaciers; A) raised marine shorelines on eastern Sjuøyane (northern Svalbard) with a post-glacial marine limit approximately 28 m a.s.l., B) polythermal, tidewater margin of Nordenskiöldbreen (inner Isfjorden) retreating back from its Late Holocene maximum, C) thin, cold-based alpine ice cap located to the northeast of Longyearbyen, central Spitsbergen, D) Wahlenbergbreen heavily crevassed and surging into Isfjorden, Aug. 2017. As a function of the size, thickness, and local climate, Svalbard glaciers exhibit a range of thermal regimes. The majority of glaciers and ice caps on Svalbard are polythermal suggesting areas of the ice are temperate (i.e., thick enough to exceed the pressure melting point) while other areas are cold-based (i.e., frozen to the substrate beneath the glacier; Björnsson et al. 1996). Some of the smallest glaciers systems are entirely cold based. These glacier systems have a low erosive impact on underlying substrate, as they do not flow across the substratum and often exhibit annual ice velocities on the mm – cm scale moving slowly due to internal deformation (Liestøl 1993; Hagen et al. 2003). Small cirque and valley glaciers characterize north central and central Spitsbergen reflecting the low precipitation there (Humlum 2002). Glacier equilibrium line altitudes (ELA) range from less than 200 m to over 700 m a.s.l. and are closely linked to precipitation patterns as well as regional summer temperatures and ice velocities (Hagen et al. 2003). In northeastern Spitsbergen and Nordaustlandet, glacier thickness has been measured and modeled to exceed 600 m (van Pelt et al. 2013; Navarro et al. 2014, 2016; Fürst et al. 2017). In numerous locations, up- stream of tidewater glacier termini, some of this grounded ice is believed to extend down to 200 meters below sea level (Fürst et al. 2017). 8 1.1.4 Surge-type glaciers, behavior and landform assemblages Svalbard has the greatest density of surge-type glaciers in the world (Hagen et al. 1993; Sevestre & Benn 2015). A glacier “surges” when it undergoes a rapid increase in velocity (order of magnitude greater than normal) and often also increases in length over a relatively short duration of a time i.e., months to a decade (Meier & Post 1969; Kamb et al. 1985; Sharp 1988; Sevestre & Benn 2015; Ingólfsson et al. 2016). A glacier that exhibits surge-type behavior typically displays these dynamic phases of ice flow, followed by a quiescent phase, a period of slow flow, and ice stagnation lasting decades to centuries (Kamb et al. 1985; Harrison & Post 2003; Dowdeswell et al. 1991). The classic model of a surge-type glacier is characterized by a regularly timed, surge event and a quiescent phase cycle. In a typical cycle a bulge of long-collected ice mass in the accumulation zone, is passed down glacier in a kinematic wave resulting in extensional stresses. The surge is followed by ice front retreat, profile steepening and slow regaining of ice mass in the accumulation zone (Clarke et al. 1984; Kamb et al. 1985; Murray et al. 2000). Traditional theory of surge-type glaciers suggests that ice front fluctuations are a result of internal dynamics and are unrelated to climatic conditions (Meier & Post 1969; Sharp 1988). These internal dynamics have been associated with the reorganization of the basal hydrological system (Kamb et al. 1985), switching of glacier thermal regime (Fowler et al. 2001; Sevestre et al. 2015), as well as sediment deformation and/or decoupling at the base of the glacier (Clarke et al. 1984; Björnsson 1998; Kjær et al. 2006). A conceptual, unifying theory regarding the controlling factors has been proposed based on the enthalpy balance of a glacier system (Sevestre & Benn 2015). Despite the theoretical strength and simplicity of this traditional surge-type glacier model, glacier surges do not always follow it. Svalbard tidewater glaciers have been observed to exhibit surge behavior that is inconsistent with the traditional model (Sevestre et al. 2018). For example, numerous marine-terminating glaciers in Svalbard have exhibited a snout destabilization where surge-type behavior initiates at the terminus and propagates upward through the glacier system (Rolstad et al. 1997; Luckman et al. 2002; Dowdeswell & Benham 2003; Murray et al. 2012; Flink et al. 2015; Dunse et al. 2015; Strozzi et al. 2017; Sevestre et al. 2018). Furthermore recent studies have highlighted a connection between surge cyclicity and mass balance (Dowdeswell et al. 1995; Striberger et al. 2011), as well as surge-type glacier distribution and climatic conditions (Sevestre & Benn 2015). These observations suggest greater complexity of surge cyclicity, dynamics and glacier front fluctuations in a warming climate (Sevestre et al. 2018; Willis et al. 2018). Glacier surges are infrequently observed, as glaciers spend the majority of their time in the quiescent phase. Thus, the identification of a surge-type glacier can be challenging. Consequently, 9 studies have developed varying approaches to identify these types of glaciers during the quiescent phase. Studies from Iceland have highlighted how surge-type glaciers respond uniquely to climatically favorable conditions (a period of positive mass balance) compared to non-surge-type glaciers (Björnsson 1998; Sigurðsson 1998; Björnsson et al. 2003). The studies conclude that the outlet glaciers known to exhibit surge-type behavior move too slowly, relative to accumulation rates, to stay in balance. In contrast, non-surge-type glaciers exhibit velocities similar to steady-state or balanced velocities (Björnsson et al. 2003). Thus, during a period of prolonged positive mass balance, typical glaciers advance, while surge-type glaciers accumulate mass and exhibit gradual steepening of surface profiles. Additionally, studies suggest that surge-type glacier behavior can be interpreted based on a distinguishable landform assemblage with specific landforms indicative of rapid ice flow (Evans & Rea 1999, 2003; Ottesen et al. 2008; Brynjólfsson et al. 2012; Schomacker et al. 2014; Flink et al. 2015; Farnsworth et al. 2016; Ingólfsson et al. 2016). Landform assemblages include streamlined features (flutes, drumlins and mega-scale glacial lineations) as well as deposits oriented oblique or perpendicular to ice flow (crevasse squeeze ridges, concertina eskers and glaciotectonized end moraines; Ingólfsson et al. 2016). This “landform assemblage” approach allows for the identification of surge behavior during the prolonged quiescent phase, as a glacier snout downwastes and upper accumulation zone regains mass. However, as geomorphological record is effective at identifying glacier systems that have exhibited rapid ice flow, it is less helpful at providing evidence for surge cyclicity (unless linked to stratigraphy; Schomacker et al. 2014). Landform assemblages similar to those of surge-type glaciers have been investigated at the margins of Greenland and Antarctic ice sheets as well as in regions covered by past ice sheets in North America, Iceland and the Barents Sea (Ó Cofaigh et al. 2002; Andreassen & Winsborrow 2009; Jakobsson et al. 2011; Dowdeswell et al. 2014; Principato et al. 2016). The combinations of streamlined and oblique landforms that have been mapped at present (and past) ice margins are associated with streaming glacier ice (Stokes et al. 2013; Spagnolo et al. 2014; Jakobsson et al. 2018). The similarities between the geomorphological assemblages of surge-type glaciers and ice streams may suggest a dynamic commonality (Kjær et al. 2008; Andreassen et al. 2014; Ingólfsson et al. 2016). The abundance and relative accessibility of surge-type glaciers on Svalbard allow for the potential to investigate analogs to large ice streams both modern and past (Flink et al. 2017). 1.1.5 Aims, and research Objectives The broad purpose of this doctoral project is to investigate the history of Svalbard glaciers through the Holocene (Fig. 3). Mapping and dating glacier marginal fluctuations allows one to estimate 10 when and where a glacier advanced or retreated. Furthermore, the comparison of palaeoclimate conditions and patterns in glacier activity may indicate the potential driving factors controlling ice marginal fluctuations. Through detailed stratigraphic and geomorphological investigations of glacier forelands, it is possible to trace the style and dynamics of (de-)glaciation. This study uses a suite of data from marine, terrestrial and lacustrine archives in order to reconstruct a mosaic of glacier and climate history for the Holocene on Svalbard. Fig. 3 Timeline of Late Pleistocene and Holocene with targeted time span for each manuscript (Cohen et al. 2013; updated). Note manuscripts from chapters II, III and V span the entire Holocene including the end of the Late Pleistocene (Farnsworth et al. 2017, in prep.; Larsen et al. 2018). The studies presented in chapters I and IV focus on the transition into the Holocene (Farnsworth et al. 2018) and the most recent period of the Holocene (Farnsworth et al. 2016). Holocene studies have long targeted the unknown aspects of the glacier and climate history of Svalbard. This doctoral thesis addresses several of these outstanding research questions: (1) Is there evidence of Younger Dryas cooling and glacier re-advances on Svalbard? (2) Did meltwater from the collapsing Laurentide Ice Sheet (LIS) influence climate on Svalbard during the 8.2 ka BP event? (3) When was the Holocene (glacial) minimum and how extensive was the glacier cover during this period? (4) When was glacier cover most extensive during the Holocene period and was the LIA the climax of the Neoglacial during the Late Holocene? (5) Which sedimentary archives provide the deepest understanding of Holocene glaciers and climate? (6) What factors control glacier fluctuations on Svalbard through the Holocene? This thesis is comprised of five chapters that detail Holocene glacier history on Svalbard. Chapters II, III and V span the entire Holocene period. Chapters I and IV target the transition from Late Pleistocene into the Holocene and glacial conditions at the end of the Holocene, respectively (Fig. 3). Chapters II and III are case studies of western Spitsbergen fjords (Van Mijenfjorden and St. Jonsfjorden) where marine and terrestrial data are used to detail glacier history. Chapters I and IV survey the entire Svalbard region targeting Early Holocene and Late Holocene glacial deposits respectively. Chapter V is a synthesis of all Holocene studies from Svalbard. 11 2. METHODS 2.0.1 Terrestrial stratigraphy and mapping Field sites were selected based on detailed investigation of aerial imagery, topographic data and digital elevation models provided by the Norwegian Polar Institute on the TopoSvalbard website. Sites were prioritized based on cross-cutting relationships of large-scale geomorphic features such as: beach ridges, glacial deposits, and drainage pathways in addition to the potential for natural geological sections cleared by hand. Lithostratigraphic logging was conducted in natural sections, and documented by sedimentological field logs, photographs, and GPS waypoints. Field mapping was conducted using computers equipped with and integrated GPS and ArcGIS software (Larsen et al. 2018) as well as aerial image prints and a hand-held Garmin GPS (Farnsworth et al. 2017, 2018). Terrestrial maps were constructed using digital ortho-rectified imagery produced by the Norwegian Polar Institute in ArcGIS 10.3. The landforms and surface sediment mapped in Van Mijenfjorden were based on 3D projected digital aerial photographs using ESRI ArcMap software with the Stereo Analyst plugin. The maps of Quaternary geology generally follow standard methodology for the Geological Survey of Norway, with some features unique to the Svalbard region (Farnsworth et al. 2017, 2018; Larsen et al. 2018). 2.0.2 Marine data acquisition and processing Sea floor data are presented from Van Mijenfjorden and St. Jonsfjorden (Farnsworth et al. 2017; Larsen et al. 2018). Data from Van Mijenfjorden was collected in 2001 by the Norwegian Hydrographic Service with an EM-1002 Simrad multibeam echo-sounder using 111 beams. The data from inner Van Mijenfjorden (Rindersbukta) were collected in July 2006 by a portable Geoswath system with a 250 kHz transducer mounted in front of a small vessel. The data from Van Mijenfjorden were gridded with a cell size of 5 m and from Rindersbukta with 1 or 5 m. Sea-floor data were collected throughout St. Jonsfjorden in June 2013. Bathymetric data were collected on board UNIS R/V ‘Viking Explorer’ with a Kongsberg EM2040 multibeam echosounder. Sea-floor sediment surface and subbottom acoustic sediment structures were studied with the Edgetech 2000 CSS combined side-scan sonar and subbottom profiler systems. The data from all surveys were tidal corrected and spikes were removed (Farnsworth et al. 2017; Larsen et al. 2018). The marine geological maps from St. Jonsfjorden were produced with QPS Fledermaus Software package based on high-resolution seafloor data. 12 2.0.3 Sampling and radiocarbon dating Following field sampling, shell, driftwood, terrestrial plant remains and whalebone subsamples were cleaned, weighed, photographed, identified and subsequently sent for dating at a laboratory with accelerator mass spectrometry (AMS). Samples presented in this thesis were radiocarbon dated at the Ångström Laboratory (Uppsala University, Sweden, Ua) the GeoBiosphere Science Centre (Lund University, Sweden, LuS), the National Laboratory for Age Determination in Trondheim (Norwegian University of Science and Technology, Tra), and at the Scottish Universities Environmental Research Centre (University of Glasgow, SUERC). See chapter V for details on previously published ages from additional laboratories (Farnsworth et al. in prep.). New and previously published radiocarbon ages are presented in tabularized form with metadata in accordance to Hughes et al. (2016; Farnsworth et al. 2017, 2018, in prep.; Larsen et al. 2018). While radiocarbon ages are given as conventional ages relative to 1950 (Stuiver & Polach 1977; Farnsworth et al. in prep.) all ages presented in text or figures are in calibrated median ages in kilo-years before present (cal. ka BP) unless otherwise noted. All terrestrial radiocarbon dates have been (re-)calibrated with IntCal13, either using the OXCAL v4.2 or the Calib Rev. 7.0.4 programs (Ramsey & Lee 2013; Reimer et al. 2013). Additionally, each 14C age of a marine organism has been (re-)corrected for a marine reservoir effect by selecting ‘MARINE13’ and inputting a ‘Delta R’ or implementing a reservoir correction. A ΔR of 70 ± 30 which corresponding to a marine reservoir of 450 years was used to correct for the marine reservoir in the first and final chapters according to Mangerud & Svendsen (2017) and Farnsworth et al. (2018; in prep.). In Chapter II a ΔR of 20 ± 30 (recommended for Svalbard; Mangerud et al. 2006) was utilized, while a reservoir age of 440 years was subtracted from marine samples in Chapter III (according to Mangerud & Gulliksen 1975; Mangerud et al. 2006). The variation in analysis is partially due to Mangerud et al. (2006) not concluding on a single approach for North Atlantic marine reservoir corrections, but rather presenting two different views on how to analyze the marine reservoir for the region. It is suggested that Atlantic Water from Scotland to Svalbard has nearly the same marine reservoir age (Mangerud et al. 2006; Mangerud & Svendsen 2017), although this may have varied through time depending on the mixture of water masses entering south of Scotland. On the Holocene timescale, variations are believed to be negligible (Bondevik et al. 1995). 13 3. AUTHOR CONTRIBUTIONS AND CHAPTER SUMMARIES All co-authors contributed to the writing of each manuscript. Farnsworth led the production of the manuscripts in Ch. I and III – V with support from co-authors. Researcher Eiliv A. Larsen from the Geological Survey of Norway led the production of the manuscript presented in Ch. II, while Farnsworth assisted with fieldwork, mapping, figure development, literature review, data interpretation and writing. All co-authors support the use of these manuscripts in this thesis. Table 1 summarizes the work conducted by the authors of each study (Chapters I – V). Task: Ch. I Ch. II Ch. III Ch. IV Ch. V Logistics / preparation Farnsworth Ingólfsson Schomacker Larsen Lyså Rubensdotter Farnsworth Ingólfsson Noormets Alexanderson Henriksen -NA -NA Fieldwork Farnsworth Ingólfsson Allaart Schomacker Larsen Lyså Rubensdotter Farnsworth Farnsworth Ingólfsson Noormets Allaart Alexanderson Henriksen -NA -NA Funding Ingólfsson Retelle Håkansson Schomacker Larsen Lyså Jensen Ingólfsson Noormets -NA -NA Mapping -NA Larsen Lyså Rubensdotter Farnsworth Farnsworth Allaart Noormets Schomacker Farnsworth -NA Figures Farnsworth Larsen Lyså Rubensdotter Farnsworth Ottesen Farnsworth Allaart Farnsworth Farnsworth Literature Review Farnsworth Larsen Lyså Rubensdotter Farnsworth Farnsworth Ingólfsson Farnsworth Ingólfsson Farnsworth Alexanderson Data interpretation Farnsworth Ingólfsson Retelle Schomacker Larsen Lyså Rubensdotter Farnsworth Nadeau Ottesen Farnsworth Ingólfsson Noormets Alexanderson Henriksen Schomacker Farnsworth Ingólfsson Retelle Schomacker Farnsworth Text prep. All authors All authors All authors All authors All authors 14 3.1 Chapter I Farnsworth, W.R., Ingólfsson, Ó., Retelle, M., Allaart, L., Håkansson, L., Schomacker, A. (2018): Svalbard glaciers re-advanced during the Pleistocene-Holocene transition. Boreas 47, 1022-1032. DOI:10.1111/bor.12326. The goal of this study was to feature the extensive and widespread moraines formed by glacier re- advances across Svalbard during the transition from the end of the Pleistocene to Early Holocene. In this manuscript we introduce the first well-dated end moraine formed during the Late-glacial Early Holocene (LGEH) in De Geerbukta, NE Spitsbergen. This landform was deposited by an outlet glacier re-advancing into a fjord extending 4.4 km beyond the Late Holocene (LH) maximum. Furthermore, we introduce six additional locations where glacier moraines have been wave-washed or cut by postglacial raised marine shorelines, suggesting the landforms were deposited before or during high relative sea-level stands, thus exhibiting a similar LGEH age. Our new evidence suggests that the LGEH glaciers were more dynamic, exhibited re-advances, and extended well beyond the extensively studied LH glacial expansion. The timing of the De Geerbukta glacier re- advance compared to a synthesis of existing data including four palaeoclimate records and 15 other proposed glacier advances from Svalbard does not suggest clear synchronicity in glacial and climatic events. The widespread occurrence of the LGEH glacier deposits on Svalbard, suggests that the culmination of the Neoglacial advances during the Little Ice Age does not mark the maximum extent of most Svalbard glaciers since regional deglaciation; it is just the most studied and most visible in the geological record. This study contributes to our understanding of the style of deglaciation during Late Pleistocene - Early Holocene on Svalbard by introducing a period in time characterized by glacier re-advances and discussing the complexities that exist between past glacier behavior and climate. 15 3.2 Chapter II Larsen, E., Lyså, A., Rubensdotter, L., Farnsworth, W.R., Jensen, M., Nadeau, M., Ottesen, D. (2018): Lateglacial and Holocene glacier activity in the Van Mijenfjorden area, western Svalbard. arktos. DOI: 10.1007/s41063-018-0042-2. The aim of this study was to detail the Late Weichselian and Holocene glacial history for the entire Van Mijenfjorden system, through sedimentological, morphological and chronological investigations of both terrestrial as well as marine archives. This study describes ice marginal glacier deposits related to re-advances spanning from Late Pleistocene to post Little Ice Age from the mouth of the fjord into the inner tributaries of the Van Mijenfjorden-system. The fjord exhibits a record of glacier fluctuations related to the episodic break-down of Late Weichselian ice in the catchment during the Late Pleistocene and Early Holocene as well as the episodic growth of re- advancing glacier systems throughout the Neoglacial, Little Ice Age and 20th century during the Late Holocene. Several interesting observations are made in this study: i) the glacier filling the entire Van Mijenfjorden system retreated during the Younger Dryas period without leaving evidence of regressive retreat, ii) numerous (Early Holocene and Neoglacial) glacier re-advances are described where ice margins extended well beyond their Little Ice Age maximum positions and, iii) the Paulabreen glacier system surged at least five times in the last 650 years, with each subsequent surge advance exhibiting a less extensive maximum than the previous, resulting in an overall decrease in mass since the early LIA. This manuscript contributes to our understanding of the deglaciation and Holocene glacier fluctuations observed from within the mouth, to the head of an entire fjord-system by detailing the complexities of glacier fluctuations and climate. This study additionally documents the diversity in the maximum position of glacier margins during the Holocene. 16 3.3 Chapter III Farnsworth, W.R., Ingólfsson, Ó., Noormets, R., Allaart, L., Alexanderson, H., Henriksen, M., Schomacker, A. (2017): Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard. Boreas 46, 585-603. DOI: 10.1111/bor.12269. The objective of this study was to reconstruct the Holocene glacial history of inner St. Jonsfjorden, western Spitsbergen, by constructing detailed geomorphological maps of marine and terrestrial environments as well as constraining the chronological sequence of glacier events. Stratigraphic and geomorphology evidence indicate an Early Holocene tributary glacier advance constrained to the transition from Late Pleistocene to Early Holocene. Identification and 14C dating of the thermophilous bivalve Modiolus modiolus to 10.0 + 0.12 cal. ka BP constrains the minimum age of deglaciation for the inner fjord, and suggests a rapid northward migration of the species during the Early Holocene. Furthermore, evidence from the Late Holocene enhances our understanding of the onset and subsequent climax of the Neoglacial-Little Ice Age in inner St. Jonsfjorden. The present- day terminus of Osbornebreen, the dominating glacier system in St. Jonsfjorden, is located over 8.5 km up-fjord of its Neoglacial maximum extent. Cross-cutting relationships suggest subsequent advances of all the smaller glaciers in the inner fjord following the break-up of the Osbornebreen tidewater glacier. Glacial deposits, landforms and their cross-cutting relationships observed in both terrestrial and marine settings imply a complex and highly dynamic environment through the later part of the Holocene. Similar to Van Mijenfjorden, much of St. Jonsfjorden becomes ice-free during the Late Pleistocene and Early Holocene. Additionally, a large dynamic re-advance from the main tidewater glacier Osbornebreen occurs during the early LIA. Since this re-advance Osbornebreen has exhibited a series of retrogressive ice marginal advances characterized by each maximum position less extensive than the subsequent. This study contributes to our understanding of the Holocene glacier history by describing evidence of dynamic glacier behavior preserved in the terrestrial and marine archives of St. Jonsfjorden. 17 3.4 Chapter IV Farnsworth, W.R., Ingólfsson, Ó., Schomacker, A., Retelle, M. (2016): Over 400 previously undocumented Svalbard surge-type glaciers identified. Geomorphology 264, 52-60. DOI: 10.1016/j.geomorph.2016.03.025. The purpose of this study was to identify glaciers that likely exhibited surge-type behavior during the culmination of the Little Ice Age by identifying crevasse squeeze ridges in an aerial image survey of glacier forelands throughout Svalbard. Crevasse squeeze ridges are landforms indicative of rapid-dynamic ice flow and are suggested to be unique to surging glacier land systems. Estimates vary greatly as to the actual percentage of surge-type glaciers in Svalbard, and consequently their distribution pattern is poorly understood. Recent (2008–2012), high-resolution aerial imagery from TopoSvalbard, provided by the Norwegian Polar Institute, was surveyed targeting all the terrestrial glacier forelands in Svalbard. Prior to our study, less than 280 individual glaciers in Svalbard had been documented to exhibit surge behavior. By using crevasse squeeze ridges as indicators of surge behavior, we have identified 431 additional glaciers that have exhibited surge-like advances. This is probably a modest value as the unique surge landforms were not visible in approximately one-third of the forelands with documented surge histories. Limits to the crevasse squeeze ridge technique are presented and potential controlling factors for crevasse squeeze ridge formation and preservation are discussed. This study contributes to our understanding of the extent of glacio-dynamic behavior experienced during the end of the LIA and early 20th century across Svalbard by surveying glacier forelands in search of a landform indicative of rapid ice-velocities and suggestive of previous surge- type behavior. 18 3.5 Chapter V Farnsworth, W.R., Ingólfsson, Ó., Alexanderson, H., Forwick, M., Noormets, R., Retelle, M., Schomacker, A. (in prep.): Holocene glacial and climate history of Svalbard - status, perspectives and challenges. We review published literature of Holocene glacier and climate history from Svalbard. This review endeavors to establish the state of the art regarding the Holocene history of Svalbard, by synthesizing findings from all Holocene studies and creating a geochronological database of all ages (14C, TCN and Lum.) spanning 12.0 ka BP to present sampled on Svalbard and the surrounding region. A database of ages has been constructed, quality assessed, and categorized by archive (marine, terrestrial and lacustrine). No review has yet been compiled despite over a century of ice front observations, hundreds of Holocene glacier studies and an ever-developing understanding of ice dynamics and the Arctic climate system. This overview: (1) presents a brief summary of major shifts in climate and glacier cover across the Svalbard region throughout the Holocene; (2) introduces a quality assessed database of published ages that constrain glacier fluctuations (deglaciation, ice free, re-advance and marginal position) and climatic conditions (warming, cooling, wetter, and drier); (3) discusses challenges in methodology as well as potentials regarding sedimentary archives and; (4) addresses the complexities of glacier systems in response to changes in climate. This synthesis establishes the state of the art regarding Holocene glacier and climate history by summarizing all Holocene literature from Svalbard and discussing findings within a modern perspective. 19 4. DISCUSSION 4.1 Svalbard Holocene climate and the unknowns While orbital conditions and the northern hemisphere summer insolation have often been linked with the Holocene thermal optimum and the subsequent Neoglacial cooling, observations of Holocene glaciers and climate from Svalbard deviate in two key ways (Laskar et al. 2004). The marine thermal optimum on Svalbard pre-dates the peak in summer insolation (Hald et al. 2007; Mangerud & Svendsen 2017). Furthermore, Svalbard glaciers are found to have exhibited marked re-advances throughout both the Early and the Late Holocene (Farnsworth et al. 2018; in prep.). These two characteristics of the Holocene climate and glacier history suggest that more than orbital forcing is controlling the temperature, precipitation and glacier behavior in Svalbard during the Holocene. In this section, the importance of Atlantic waters around Svalbard with regard to sea-ice cover, temperature, and precipitation is discussed further. 4.1.1 Atlantic waters control the temperature Today, mean annual air temperatures recorded on Svalbard (Longyearbyen) are at least 5 °C higher than other high-Arctic meteorological stations (Eckerstorfer & Christiansen 2011). Svalbard’s mild, high-Arctic climate is a direct result of the oceanic currents arriving to and traveling around the archipelago today (Førland et al. 1997; Drange et al. 2013). Furthermore, current sea-ice cover conditions (i.e., ice-free conditions on the west coast and the sea-ice dominated east coast) drive a strong temperature gradient across Svalbard (Isaksen et al. 2016). Thus, the West Spitsbergen Current sea surface temperature anomalies influence regional sea-ice extent which, in turn, impacts regional and local variations in air temperature and precipitation (Jung et al. 2017). Svalbard’s Holocene climate, and by association its glaciers, are linked to the effects of North Atlantic waters arriving to the region. Further modeling investigations are needed to address the controls and robustness of the West Spitsbergen Current. This will improve knowledge of the over-arching factors governing Holocene climate on Svalbard. 4.1.2 Palaeo-precipitation remains unknown There is no direct means of measuring precipitation during the majority of the Holocene on Svalbard. The oldest ice core records showing snow accumulation extend back to roughly 1 ka BP (Isaksson et al. 2005; Divine et al. 2011). The greatest potential for reconstructing Holocene precipitation on Svalbard is through detailed investigations of lake catchments and sedimentary archives (Røthe et al. 2018). The first hydrogen isotope leaf wax study from Svalbard has recently 20 been conducted on a Holocene lake record and provides insight in to past hydroclimate (Balascio et al. 2018). The analysis of leaf wax hydrogen isotopes from lake sediments is a developing proxy associated with hydroclimate, palaeo-precipitation and sedimentation within lake catchments (Thomas et al. 2012, 2016). As this method further develops and studies target lakes across Svalbard, our understanding of past precipitation across the region will improve (Farnsworth et al. in prep.). Although precipitation patterns on Svalbard through the Holocene are not well understood, the knowledge of modern processes can help contextualize Holocene precipitation (Humlum 2002). Thus, the better the understanding of the intricacies of modern precipitation processes, more effective models will be developed for both past and future conditions. Numerous studies of modern field observations and modeling suggest interconnection between increased temperature, decreasing sea-ice extent, and increased precipitation for the Norwegian High Arctic (Nowak & Hodson 2013; Bitanja & Selten 2014; Isaksen et al. 2016; Kopec et al. 2016). Increased rates of precipitation recorded in Svalbard have been associated with warmer autumn seasons and the associated delay of sea-ice formation (Christiansen et al. 2013; Nowak & Hodson 2013). Relatively warm and wet conditions have driven strong negative mass balance on low elevation cirque and valley glaciers (Möller & Kohler 2018). Alternatively, continental high-elevation glacier accumulation zones, which remain above the freezing point, may gain mass and benefit from the increased precipitation. Consequently, precipitation during the Early Holocene, due to the rapid increase of ocean temperatures (Hald et al. 2004), and decrease in sea-ice cover (Müller & Stein 2012), likely had a large, but as of yet unquantifiable influence on the regional hydro-climate and glacier mass balance (Mangerud & Svendsen 2017; Farnsworth et al. 2018; Balascio et al. 2018). 4.2 Drivers of Holocene glacier re-advance Svalbard glaciers re-advanced throughout the Early Holocene (as late as 9 ka BP) and through the entire Late Holocene (Larsen et al. 2018; Farnsworth et al. 2017, 2018, in prep.). However, the rate of glacier growth during these periods of re-advances is poorly constrained and not fully understood (Farnsworth et al. 2018). There is currently no evidence to suggest that glaciers re-advanced through the Mid Holocene on Svalbard (Farnsworth et al. in prep.). 4.2.1 Early Holocene re-advances One must consider what evidence suggests that Early Holocene glacier re-advances are controlled by mass balance and hence are climatically forced. Ice-marginal oscillations during the Early Holocene are asynchronous, and appear to occur during an unfavorable climate (warm oceans, high 21 summer insolation, variable sea-ice conditions; Larsen et al. 2018; Farnsworth et al. 2017, 2018, in prep.). Although the percentage of glacier cover across Svalbard may have theoretically increased at periods during the deglaciation (due to the re-advancing glacier margins), general ice volume was probably in continuous decline during this period (Farnsworth et al. 2018; in prep.). Unless increased precipitation during the Early Holocene out-weighed the present reconstructed climatic factors, ice marginal re-advances were probably controlled by glacio-dynamic behavior. 4.2.2 Neoglacial – Little Ice Age re-advances Svalbard glaciers re-advanced throughout the Neoglacial period (Werner 1993; Reusche et al. 2014; Røthe et al. 2015; van der Bilt et al. 2016; Philipps et al. 2017) and during the LIA (Svendsen & Mangerud 1997; Snyder et al. 2000; Humlum et al. 2005; de Wet et al. 2018). The combination of low summer insolation (Laskar et al. 2004) and explosive volcanic activity (Miller et al. 2012) influenced cooling in the northern hemisphere and may have been key factors in creating conditions favorable for glacier growth during the Late Holocene. As introduced in the last section, little is known about patterns of Holocene precipitation across Svalbard. One alkenone-based (summer) temperature reconstruction from western Spitsbergen lake sediments suggests the Little Ice Age was “mild” (D’Andrea et al. 2013). The authors go on to suggest that precipitation played a larger contribution to regional glacier re-advances than previously acknowledged (D’Andrea et al. 2013). While it is important to note the limited knowledge of past precipitation, there is ample evidence suggesting that air and ocean temperatures were relatively cool and also favored glaciers growth during the late Neoglacial and LIA (Divine et al. 2011; Bartels et al. 2017; Ojala et al. 2018; Røthe et al. 2018; van der Bilt et al. 2018; Balascio et al. 2018). The Kongressvatnet lake sediment record suggests that, in the 18th and 19th centuries, reconstructed summer temperatures were 2 °C to 3 °C cooler during the later LIA by comparison to the 20th century. Based on the 100-year-plus temperature record from Longyearbyen, an increase of 1.5 °C in summer temperatures have had a substantial influence on glaciers. Furthermore, annual variability reflected in the record’s mean annual air temperature (seen on an annual-decadal scale) is dominantly derived from fluctuations during the winter and shoulder seasons (Christiansen et al. 2013; Farnsworth et al. in prep.). Consequently, it is not clear how much impact 2 – 3 °C (average summer) temperature increase would have on Svalbard glaciers. However, it is roughly twice the increase recorded during the last century in Longyearbyen. The reconstructed summer temperatures from the Kongressvatnet lake sediment record also suggest the Neoglacial and early LIA was even cooler than the late LIA. Accordingly, most Late Holocene glacier re-advances (dated by overridden vegetation or mollusc shells) date to the early (between 1.0 – 0.5 ka BP), not the late LIA (Farnsworth et al. in prep.). Early observations from the 22 1900s, describing glaciers proximal to their Late Holocene maxima does not suggest anything about the timing of glacier re-advance. It just suggests the glaciers were slowly starting to retreat. The D’Andrea et al. (2013) summer temperature record is particularly interesting as it does not reflect a clear or defined cooling excursion during the LIA, but rather an extended and pronounced phase of Neoglacial. Furthermore, the study raises the question, are average summer temperatures sensitive “first responders” to climatic shifts in Svalbard? Sea-ice conditions during the Neoglacial and LIA are characterized by increasing but variable ice extent around Svalbard (Müller & Stein 2012; Müller et al. 2014; Bartels et al. 2017). Although increasing sea-ice cover reconstructed through most of the Late Holocene is presumed to favor glacier growth (suppressing summer temperatures, minimizing the number of positive degree days and decreasing frontal ablation) it is unclear how much this restricts precipitation (Farnsworth et al. in prep.). In phase with climatic favorability, several glacier re-advances constrained to the Neoglacial – LIA have been characterized as surges based on size, extent of glacial deposits and preservation of landforms (related to rapid ice advances) corresponding to associated ice-margins (Ottesen et al. 2008; Kristensen et al. 2009; Kempf et al. 2013; Farnsworth et al. 2016, 2017; Lovell & Boston 2017; Flink et al. 2017; Lyså et al. 2018). While the most extensive Late Holocene glacier deposits have been associated with surge-type behavior especially near the culmination of the LIA (Kristensen et al. 2009; Kempf et al. 2013; Flink et al. 2015; Lyså et al. 2018), an increasing number of studies have identified both complete and fragmented moraine ridges outboard of the LIA maxima (Werner 1993; Sletten et al. 2001; Reusche et al. 2014; Philipps et al. 2017; Larsen et al. 2018). The classical perspective of the (Late) Holocene glacial maximum occurring during the culmination of the LIA is being challenged (Svendsen & Mangerud 1997; Snyder et al. 2000). 4.2.3 Post Little Ice Age re-advances Despite trends in negative mass balance throughout the later 20th century (Hagen et al. 2003; Nuth et al. 2013; Martin-Moreno et al. 2017; Østby et al. 2017) glacier re-advances have been observed after the LIA. However, these glacier re-advances are not sustained in the long-term and are attributed to short-lived changes in glacio-dynamic conditions such as surging (Lefauconnier & Hagen 1991; Hagen et al. 1993, 2003; Dunse et al. 2015; Lovell et al. 2015; Sevestre & Benn 2015; Sevestre et al. 2018). 23 4.2.4 Early vs. Late Holocene glacier cover variation Although the percent of glacier cover from the LIA through the 20th century may be similar to the percentage during the Early Holocene (c. 10.0 ka BP), the spatial distribution of glacier ice across Svalbard is different (Fig. 4). At numerous sites on Svalbard, Late Holocene glaciers have re- advanced over shorelines of Late Pleistocene-Early Holocene age (Fig. 4A, C & E). Fig. 4 The distribution of glacier-cover across Svalbard is different during the Early Holocene compared to the Late Holocene. A, C & E) Examples of aerial imagery and Landsat data where Late Holocene glaciers overlay or have overridden shorelines of Late Pleistocene and Early Holocene age; A) Scottbreen (Mangerud & Landvik 2007), B) Storøya, (Jonsson 1983) and Augnebreen (Bondevik et al. 1995). B, D & F) Examples of topographic maps from selected regions 24 with previously described Early Holocene ice margins. At present, catchments are ice-free with no indications of Late Holocene glaciation in; B) Finn Malgrenfjorden (Ulfstedt 1987), D) Gipsdalen (Tolgensbakk 1990) and F) Lyckholmdalen (Farnsworth et al. 2018). In some cases, raised beaches from high relative sea levels (of Early Holocene age) are seen melting out from cold-based ice margins (Fig. 4C). Evidence from these sites implies that these areas were ice-free at the time of shoreline formation and that subsequent climatic changes effectively reinvigorated glaciers during the Late Holocene. In contrast, there are examples of unglaciated regions today where Early Holocene ice margins re-advanced, suggesting that the residual SBSIS- ice persisted in these regions and subsequent Holocene climate was not conducive to re-glaciate these catchments. The present general distribution of glaciers ice and prevailing winds suggests that these regions are dry and shadowed from precipitation (Fig. 4; Humlum 2002). It is unclear if these variations in ice cover relate to dynamic behavior during the Early Holocene, trends in Late Holocene precipitation patterns or some combination of the two. 4.3 Advances in Deglaciation Through the 20th century, Svalbard glaciers have undergone widespread retreat (Blaszczyk et al. 2009; Nuth et al. 2013; Martin-Moreano et al. 2017). Air temperatures have risen (Christiansen et al. 2013; Isaksen et al. 2016); sea-ice cover has declined (Muckenhuber et al. 2016; Petty et al. 2018); and ocean temperatures have increased, as evidenced by the reappearance of thermophilous marine molluscs around Svalbard (Berge et al. 2005; Mangerud & Svendsen 2017). Despite these trends, some Svalbard glaciers continue to re-advance. Current conditions, regarding atmospheric and ocean temperatures, are similar to those reconstructed from the period of glacier retreat during the Early Holocene (Hald et al. 2004; Muller & Stein 2012; Hormes et al. 2013; Bartels et al. 2017; Mangerud & Svendsen 2017). However, there are two clear contrasts between these periods generally characterized by glacier retreat. First, northern hemisphere summer insolation approached its maximum during the Early Holocene (Fig. 4; Laskar et al. 2004). Second, post-glacial isostatic rebound rates were at their peak during the Early Holocene, while, at present, sea level in most regions on Svalbard has stabilized or is transgressing (Bondevik et al. 1995; Forman et al. 2004; Fjeldskaar et al. 2018). In the following section several hypothetical scenarios, which may have influenced glacier dynamics during the Early Holocene on Svalbard, are discussed. 4.3.1 Restrained rebound and ice caps chasing the equilibrium line Relative sea level curves from Svalbard suggest that half of the total glacio-isostatic uplift occurs during the first 2 – 2.5 ka after deglaciation (Landvik et al. 1998; Forman et al. 2004). Unlike post- 25 glacial relative sea level curves constructed from dated raised marine shorelines, the rate and extent of uplift prior to local deglaciation is invisible in our raised marine records (Fig. 5A). The majority of the sea level curves from Svalbard do not project “restrained rebound” as described by Andrews (1970), but rather rapid post-glacial uplift (Bondevik et al. 1995; Forman et al. 2004). While raised shorelines provide visual geomorphological evidence of postglacial uplift, restrained rebound is the term used to describe pre-deglacial rebound. Thus, some extent of uplift presumably initiated prior to the majority of Svalbard shoreline formation (coastlines becoming ice-free). This section speculates about the amount of (restrained) rebound that occurred during Svalbard’s deglaciation prior to the formation of shorelines, and the start of the geomorphological record. Fig. 5 Introduction and estimates of pre-deglacial (restrained) rebound. (A) Schematic relative sea level curve from Svalbard with the inverse relative uplift of 100 m after shorelines became ice-free slightly before 11 ka BP (Forman et al. 2004). (B) Model of the distribution ratio of ice sheet situated above to below sea level once isostatic depression reaches equilibrium (modified from Wolcott 1970; Andrews 1974). (C) Estimates of the extent of restrained rebound prior to an 11 ka BP deglaciation based on an ice sheet thickness of 2100 m and varying levels of isostatic equilibrium. (D) Minimum estimate of restrained rebound (1/6th of equilibrium = 100 m) projecting a schematic ice cap traveling through a constant palaeo-equilibrium line altitude. 26 Assuming ice thickness around Svalbard and the Barents Sea was on the order of 2000 m thick, at equilibrium, a ratio of 2:1 would reflect the distribution of ice above to ice below sea level (Andrews 1974; Lambeck, 1995; Landvik et al. 1998; Fig. 5B). Additionally, Svalbard is one of the regions glaciated longest throughout the last glacial cycle (~18 ka; Patton et al. 2015). If the ice mass persisted for long enough to arrive at equilibrium, theory would suggest roughly 700 m of the ice was located below sea level, while 1400 m of it would remain above the relative sea level. Given the assumptions of the thickness, duration and thinning of the ice sheet around Svalbard, it is only possible to speculate if equilibrium was reached regarding ice load over the (20 km) thick continental crust (Ritzmann & Jokat 2003; Patton et al. 2015). The SBSIS lost the bulk of its mass between 18 – 14 ka BP (Fig. 1; Hughes et al. 2016). While thinning may have initiated early in southwestern Spitsbergen, the Svalbard region held ice later than the Barents Sea, with the majority of the ice loss occurring between 15 – 10 ka BP (Fig. 1C; Hughes et al. 2016; Young et al. 2018). Various schematic scenarios of linear uplift (restrained rebound ) are projected between 15 – 11 ka BP based on an ice thickness of 2100 m and varying levels of isostatic equilibrium (Fig. 5C). These projections are linked to a schematic relative sea level curve with 100 meters of post-glacial isostatic uplift between 11.2 ka BP and present. Isostatic adjustment presumably reached at least 1/6th of equilibrium, thus there was a minimum of a 100 m of hidden uplift between 15 – 11 ka BP, prior to shorelines becoming ice free. Similar to Ruddiman’s (2006) schematic model of glaciation, isostatic adjustment and equilibrium line altitude (ELA), the model focuses on the interval of deglaciation (Fig. 5D). Although neither deglaciation nor uplift was linear, a model ice cap is plotted uplifting through a constant theoretical equilibrium line altitude (ELA). This schematic diagram represents what could have occurred during the period of pre-deglacial rebound, prior to Svalbard shorelines becoming ice free. The minimum projected restrained rebound (100 m) is nearly double the rise in ELA (but in the opposite direction) estimated for the 20th century (1900 – present; Røthe et al. 2015). The mass loss exhibited during the post-LIA period on Svalbard suggests that a 60 m rise in ELA can have a profound influence on glacier extent and mass balance. Consequently, if the relative ELA descended roughly 100 m (or if land rose up through an ELA), glaciers would exhibit strong positive mass balance. Furthermore, depending on the elevation of the ELA and the regional terrain, even a small shift in ELA can have substantial impact. Regions characterized by highland plateaus can undergo near instantaneous glaciation if terrain surpasses the threshold of the ELA (Ives et al. 1975; Lie et al. 2003). Although there are broad assumptions regarding these estimates of pre-deglacial, restrained rebound (ice thickness, duration, age of deglaciation etc.), it is possible that hidden 27 geomorphological uplift may have impacted Svalbard glaciers and ice dynamics during the deglaciation. The regions characterized by high-elevation plateaus, such as northeastern Spitsbergen and Nordaustlandet, could have benefited the most from this pre-deglacial rebound. Although the rate of uplift is only a fraction of uplift projected from records in Iceland during deglaciation (Norðdahl & Ingólfsson 2016) this process, coupled with enhanced precipitation patterns maybe one of the contributing factors influencing glacial re-advances during the Late Pleistocene and Early Holocene (Farnsworth et al. 2018). 4.3.2 Late Pleistocene – Early Holocene ice shelves on Svalbard As the WAIS provides a modern analogue for what conditions might have been like around Svalbard and the Barents Sea during the transition from Late Pleistocene to Early Holocene (Mercer et al. 1970; Esteves et al. 2017), it is important to consider how an ice shelf may have influenced regional climate and glaciers. Ice shelves or palaeocrystic ice leave subtle morphological and sedimentological fingerprints (Hjort et al. 2001; Roberts et al. 2008; Bradly & England 2008; Fitzsimmon et al. 2012; Davies et al. 2017; Furze et al. 2017). Therefore relatively few palaeo-ice shelves have been described (Hodgson 1994; England et al. 2009; Furze et al. 2018). No evidence for Holocene ice shelves has yet been identified on Svalbard. Nor, at this point, have ice shelves been the focus of any Svalbard studies. However, as ice shelves are currently an important component of the WAIS, it is possible that floating ice shelves or tongues existed during the transition from the Late Pleistocene and Early Holocene (Farnsworth et al. in prep.). The implications of the presence of ice shelves and their potential destabilization can have substantial consequences for glacier mass balance and dynamics, related to buttressing, back-stress, as well as precipitation (Cook & Vaughan 2010; Davies et al. 2017). Where ice shelves in the western Antarctic have thinned, retreated, and collapsed (Cook & Vaughan 2010), tributary glaciers have exhibited accelerated ice discharge and dynamic unsustainable advances reflective of surge-type behavior (De Angelis & Skvarca 2003; Scambos et al. 2003; Rignot et al. 2004; Berthier et al. 2012). Recent studies from Svalbard have speculated that if ice shelves could have developed from a concentrated iceberg mélange or paleocrystic ice in Van Mijenfjorden and Storfjorden during deglaciation, and influenced Early Holocene glacier dynamics (Larsen et al. 2018; Nielsen & Rasmussen 2018). Despite the mouths of Svalbard fjords becoming ice-free prior to the onset of the Younger Dryas, there are minimal ice marginal deposits identified from this period. Furthermore, marine sedimentation rates increase only at the start of the Early Holocene (Forwick & Vorren 2010; Larsen et al. 2018; Nielsen & Rasmussen 2018). More work is needed in these regions that 28 targets sediment archives spanning the Late Pleistocene to Early Holocene. Studies should investigate the geomorphological fingerprints related to ice shelves, like epi-shelf lake deposits and ice-shelf moraines (Hjort et al. 2001; Bentley et al. 2005; Furze et al. 2018). 4.3.3 Modern glacier behavior as a model for past glacier dynamics Through the 20th century, Arctic glaciers and sea-ice extent have been in a general decline (Bintanja & Selten 2014; Huss and Hock 2015). At present, Arctic glacier ice surface velocities are generally increasing (Strozzi et al. 2017) resulting in increased rates of frontal ablation (Luckman et al. 2016). Since the onset of the 20th century, the rise in both atmospheric and oceanic temperatures is observed from around the Svalbard archipelago (Pavlov et al. 2013; Isaksen et al. 2016). Frontal ablation of Svalbard tidewater glaciers is closely linked to fjord water temperatures at depth which drive ice-front undercutting by oceanic melt (Luckman et al. 2016; Vallot et al. 2018). This process is apparent from enhanced surficial crevassing (Luckman et al. 2016). Crevasses can greatly influence the glacier hydrological system allowing for increased passage ways for supra-glacial run- off to enter the glacier system, and increasing the probability of meltwater lubricating the glacier bed (Dunse et al. 2015; How et al. 2017). Recently, several tidewater glaciers in Svalbard have exhibited a snout destabilization, when surge-type behavior initiates at the terminus and propagates upward through the glacier system (Dunse et al. 2015; Strozzi et al. 2017; Sevestre et al. 2018). This development in process understanding complicates our over-simplified view of surge dynamics and glacier front fluctuations in a warming climate (Sevestre et al. 2018). Traditional surge-type glacier theory suggests that ice front fluctuations (i.e., surge events) are a result of internal dynamics and are unrelated to climatic conditions (Meier & Post 1969; Sharp 1988). However, other studies have highlighted a connection between surge cyclicity and mass balance (Dowdeswell et al. 1995; Striberger et al. 2011), as well as surge-type glacier distribution and climatic conditions (Sevestre & Benn 2015). Furthermore, glaciers can exhibit surge-type behavior both in and out of phase with climate favorable for glacier growth, given the correct mass balance and basal conditions (Dowdeswell et al. 1995; Striberger et al. 2011; Philipps et al. 2017). Svalbard glaciers re-advanced during deglaciation in the Early Holocene as well as through the 20th century, and both periods have been generally characterized as climatically unfavorable for glaciers (Farnsworth et al. 2018, in prep.). Surge-type behavior has been suggested for numerous Late Holocene glacier re-advances on Svalbard, based on (1) the extent of the glacier deposits (Liestøl 1977; Kristensen et al. 2009), (2) the landforms (assemblages) associated with rapid ice flow (Schomacker et al. 2014; Copland et al. 2011; Paul 2015; Farnsworth et al. 2016; Lovell & 29 Boston 2017) and (3) the internal deformation structures preserved within residual dead-ice (Lovell et al. 2015). However, these indicators of surge behavior are not always preserved from older events due to the erosion of glacier deposits (Kirkbride & Winkler 2012; Landvik et al. 2014). Given the absence of knowledge about past precipitation patterns, unknown rates of hidden uplift, the fingerprints of ice shelves, and our poor constraint on the chronology of Holocene glacier re-advances, we are unable to positively conclude whether climate or glacier dynamics is the driving factor of Holocene ice-front fluctuations (Farnsworth et al. 2018, in prep). However, on a Holocene time-scale, the difference of a surge-type glacier’s response to climate compared to a non- surge-type glacier maybe negligible as the assumed variability fits within the margins of error of our current dating techniques (Philipps et al. 2017). It is critical to recognize how the style of subsequent ice loss following the glacier re- advances may vary given the two different types of glacier behavior. Oerlemans & van Pelt (2015) suggest one of the key results of a surge is lowering the mean surface elevation, which increases the ablation area and subsequently causes a negative perturbation in the mass budget of the glacier system. Thus, a glacier that surges into a warming climate will exhibit a greater rate of retreat in comparison to a non-surge-type glacier (retreating under the same unfavorable climatic forcing; Dowdeswell et al. 1995; Oerlemans & van Pelt 2015; Farnsworth et al. 2018). The process of a dynamic unsustainable re-advance into a warming climate is an efficient means of shedding a large mass of ice in a short period of time. Potentially, this could result in a retreat so extensive the glacier system is not able to return to an active surge phase (Dowdeswell et al. 1995; Striberger et al. 2011). For most Svalbard’s glaciers where numerous surges have been documented, each successive surge is less extensive than the previous (Fig. 6A & B; Flink et al. 2015; Lønne 2016; Farnsworth et al. 2017; Flink et al. 2017; Larsen et al. 2018), a documented exception to this is Nathorstbreen (Sund et al. 2014). These recessive surge moraines have been mapped and described across Svalbard with the most frequent series of events identified in front of Tunabreen and Paulabreen (Fig. 6A & B; Flink et al. 2015; Larsen et al. 2018). Eight separate surge events have been documented in front of both glacier systems (four each) where submarine moraine ridges indicated the outer-most ice extent for each event (Flink et al. 2015; Larsen et al. 2018). Additionally, recent surges have been observed at both glacier systems where submarine morphology is yet to be investigated, but ice extent was less extensive than the four earlier events (Kristensen & Benn 2012; Larsen et al. 2018; St. Andrews Glaciology and C. Borstad personal comm.). In some locations on Svalbard, where postglacial uplift has raised once-shallow fjords above the modern coast, similar ice marginal deposits can be identified in series (Fig. 6C & D). 30 Given the relation between the high relative sea level and the glacier deposits, the timing of the formation of these landforms must have been rapid and shortly after deglaciation (Farnsworth et al. 2018). Fig. 6A & B) Topographic maps of documented end moraines from surges indicating retrogressive surge extent for Tunabreen and Paulabreen, respectively (Flink et al. 2015; Larsen et al. 2018). C) Topographic map with a series of ice marginal deposits preserved in Mosseldalen that were deposited successively into a high relative sea level during the Late Pleistocene and Early Holocene. Landforms from all three sites seem to have been deposited into a shallow fjord environment during a period characterized by a transition from cool to warm climate. D) Inset aerial image of Mosseldalen with ice marginal deposits indicated modified from TopoSvalbard. Model projections of ELA for the rest of the 20th century suggest ELAs will rise above most of the summits and ice fields in Svalbard (Førland et al. 2011; Østby et al. 2017). This will lead to considerable mass change as well as shifts in the glacier thermal and hydrological regimes, directly influencing glacier dynamics (Dunse et al. 2015; Østby et al. 2017; Willis et al. 2018). Given the processes detailed in recent Svalbard surge studies, it appears that warming oceans could continue to trigger instabilities at tidewater margins (Dunse et al. 2015; Strozzi et al. 2017; Sevestre et al. 2018; Willis et al. 2018). Considering the climatic similarities between the transitions of the Younger Dryas to Early Holocene and the LIA to the 20th century, it is possible that these processes 31 have acted on Svalbard glaciers in the past. The deglaciation at the onset of the Holocene may better reflect current ice dynamics than previously acknowledged. These findings from Svalbard compliment earlier studies documenting the dynamic and complex pattern of deglaciation, which appears time transgressive across the Barents Sea (Winsborrow et al. 2010; Andreassen et al. 2014; Bjarnadóttir et al. 2014; Newton & Huuse 2017). As comparisons are made between the WAIS and the SBSIS, it is increasingly important to consider unsustainable glacio-dynamic behavior regarding the style of ice loss in a warming climate (Mercer 1970; Oerlemans & van Pelt 2015; Newton & Huuse 2017; Willis et al. 2018; Farnsworth et al. 2018). Evidence of this dynamic style of deglaciation is emerging from the Svalbard Barents Sea region, during both the last deglaciation and at present. Are we beginning to see this glacio- dynamic behavior in the West Antarctic (De Angelis & Skvarca 2003; Scambos et al. 2003; Berthier et al. 2012)? 32 33 5. SUMMARY AND CONCLUSIONS Through the Holocene, Svalbard glaciers have exhibited (at least) two phases of widespread glacier re-advance, during the Early Holocene and throughout the entire Late Holocene. No geomorphological features have been identified corresponding to glacier re-advances between 9.0 – 4.5 ka BP. Consequently, relatively little is known about the extent of Svalbard glaciers during the Mid Holocene. Early Holocene glacier re-advances are identified across Svalbard, corresponding to a diverse range of glacier sizes. With our current level of age constraint, these ice marginal fluctuations do not appear synchronous in time. Furthermore, the Early Holocene climate is believed to have been warm, unfavorable for glacier growth and characterized by deglaciation. Early Holocene glacier re-advances appear to relate to the time-transgressive nature of deglaciation. Glacier re-advances likely corresponded to glacier dynamics (not mass balance) and reflect the complex style of ice-mass-loss during a changing climate. Landforms and deposits from glaciers re-advancing during the Late Holocene have been the primary focus of Holocene glacial studies. Glacier re-advances and corresponding deposits have been attributed to episodic Neoglacial cooling and the Little Ice Age (LIA). The majority of Late Holocene glacier re-advances have been dated to between 4.0 – 0.5 ka BP. The highest frequency of re-advances are constrained to 1.0 – 0.5 ka BP, during the first half of the LIA. Glacial landforms and deposits from LIA re-advances have been suggested to indicate rapid and dynamic glacier behavior (in some cases characteristic of surge-type events). During the 20th century (post-LIA), Svalbard glaciers have exhibited widespread negative mass balance, ice marginal retreat, and glacier thinning. This phase of retreat has had a direct influence on glacier thermal regime, hydrologic system and surface profile. Through the 20th century, some Svalbard glaciers have continued to exhibit surge-type re-advances. Several glaciers have exhibited this behavior numerous times. These glacio-dynamic re-advances have been un- sustained and each subsequent surge has been less extensive then the previous. Consequently, and despite re-advance, glaciers reflect a continual phase of ice-mass-loss in a periodic fashion. • Throughout the Holocene, Svalbard glaciers have responded to a varying combination of climatic, environmental and dynamic driving factors which influence both the extent and behavior of ice margins. 34 • Glaciers during the Late Pleistocene and Early Holocene were dynamic, exhibited re- advances and extended well beyond Late Holocene glacier maxima in many locations across Svalbard. • The marine Holocene thermal optimum on Svalbard, marked by the early arrival of warm water species to the coasts of Svalbard during the onset of the Holocene, pre-dates the peak in northern hemisphere summer insolation and the terrestrial thermal optimum. • During the Holocene glacial minimum, glaciers covered a small, but unknown percentage of Svalbard. Marine sediment cores from Svalbard fjords suggest some tidewater glacier fronts persisted through the glacial minimum. • Evidence of episodic Neoglacial glacier re-advances is being identified more commonly across Svalbard, suggesting a stepped and variable phase of conditions favorable for glacier growth during the Late Holocene. • The Little Ice Age maximum does not reflect the glacial maximum extent across Svalbard; since the deglaciation, the onset of the Holocene or even during the Late Holocene, in most locations. • Svalbard glaciers exhibited widespread glacio-dynamic re-advances at some point during the Little Ice Age, in some cases developing (and preserving) landforms indicative of surge-type behavior. • Following the Little Ice Age, the majority of Svalbard glaciers have undergone extensive retreat and thinning which has led to changes in the thermal regime, hydrological system, and surface profile of glaciers. • Some glaciers on Svalbard have continued to exhibit dynamic re-advances characterized as surges-type behavior, however each surge-event is less extensive than the previous exhibiting a trend of continual ice loss. 35 6. FUTURE PROSPECTS All studies of glacier and climate history should strive to have a holistic approach to systems and processes; taking into account a mixture of methods, archives and disciplines. Future Svalbard studies should endeavor to: • Enhance our understanding of the range of variability as well as factors governing the arrival and strength of the West Spitsbergen Current to Svalbard. As the impact of these warm Atlantic waters around Svalbard has a strong control on regional temperatures and precipitation, these conditions are critical for predicting future climatic conditions as well as glacier extent and dynamics. • Better address past precipitation patterns both through proxy studies (e.g., leaf wax hydrogen isotopes) and better understanding recent shifts and current processes in order to constrain models used to project precipitation in a changing climate. • Target lakes corresponding to larger glacier systems which persisted throughout more of the (or the entire) Holocene. This would close the gap in knowledge regarding the Holocene glacier minimum and the early onset of the Neoglacial. Furthermore, defining a Holocene tephrochronology for Svalbard, based on pan-Arctic volcanism, would improve high-Arctic lake sediment chronologies as well as put volcanic events into climatic perspective. • Focus on regions around Svalbard (e.g. Van Mijenfjorden, Storfjorden, Isfjorden and Hinlopen Strait) where ice shelves or palaeocristic ice may have formed during the transition from the Late Pleistocene to the Early Holocene. By combining stratigraphic and morphological data from marine and terrestrial archives it may be possible to suggest if these features existed around Svalbard and whether they had an influence on glacier dynamics or environmental conditions. • Further constrain the timing of Late Pleistocene – Early Holocene glacier re-advances. Additionally, attempts to model this transitional period based on estimates of oceanic/atmospheric temperatures, precipitation and relative uplift would be valuable. • Improve our understanding of post-glacial (and pre-deglacial) uplift around Svalbard. Studies should continue to detail sea level curves to better resolve Holocene transgressions and rebound intricacies. Further studying crustal rheology and characteristics of the response of the Earth’s crust to isostatic loading and unloading (return rate/size of area influenced) would be valuable. 36 References Andersson, T., Forman, S.L., Ingólfsson, Ó., Manley, W.F. (1999): Late Quaternary environmental history of central Prins Karls Forland, western Svalbard. Boreas 28, 292–307. Andersson, T., Forman, S.L., Ingólfsson, Ó., Manley, W.F. (2000): Stratigraphic and morphologic constraints on the Weichselian glacial history of northern Prins Karls Forland, Western Svalbard. Geografiska Annaler Series A 82, 455–470. 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(2015): Evolution of surface velocities and ice discharge of Larsen B outlet glaciers from 1995 to 2013. The Cryosphere 9, 957–969. Young, N.E., Lamp, J., Koffman, T., Briner, J.P., Schaefer, J., Gjermundsen, E.F., Linge, H., Zimmerman, S., Guilderson, T.P., Fabel, D., Hormes, A. (2018): Deglaciation of coastal south-western Spitsbergen dated with in situ cosmogenic 10Be and 14C measurements. Journal of Quaternary Science 33. DOI: 10.1002/jqs.3058. Østby, T.I., Schuler, T.V., Hagen, J.O., Hock, R., Kohler, J., Reijmer, C.H. (2017): Diagnosing the decline in climatic mass balance of glaciers in Svalbard over 1957–2014. The Cryosphere 11, 191–215. Research papers Chapter I Farnsworth et al. 2018 Chapter II Larsen et al. 2018 Chapter III Farnsworth et al. 2017 Chapter IV Farnsworth et al. 2016 Chapter V Farnsworth et al. in prep. Chapter I Svalbard glaciers re-advanced during the Pleistocene-Holocene transition Farnsworth, W.R., Ingólfsson, Ó., Retelle, M., Allaart, L., Håkansson, L., Schomacker, A. (2018): Svalbard glaciers re-advanced during the Pleistocene-Holocene transition. Boreas 47, 1022-1032. DOI:10.1111/bor.12326. Glaciotectonized shallow marine sediments on the northern side of the Fakse moraine, De Geerbukta, northeastern Spitsbergen (photo from Ingólfsson). Svalbard glaciers re-advanced during the Pleistocene–Holocene transition WESLEY R. FARNSWORTH , OLAFUR INGOLFSSON, MICHAEL RETELLE, LIS ALLAART, LENA M. HAKANSSON AND ANDERS SCHOMACKER Farnsworth,W.R., Ingolfsson, O.,Retelle,M.,Allaart,L.,Hakansson,L.M.&Schomacker,A.:Svalbardglaciers re- advanced during the Pleistocene–Holocene transition. Boreas. https://doi.org/10.1111/bor.12326. ISSN 0300-9483. Despite warming regional conditions and our general understanding of the deglaciation, a variety of data suggest glaciers re-advanced onSvalbard during theLateglacial–earlyHolocene (LGEH).Wepresent the firstwell-dated end moraine formed during the LGEH in De Geerbukta, NE Spitsbergen. This landform was deposited by an outlet glacier re-advancing into a fjord extending 4.4 km beyond the late Holocene (LH)maximum.Comparing the timing of the De Geerbukta glacier re-advance to a synthesis of existing data including four palaeoclimate records and 15 otherproposedglacier advances fromSvalbarddoesnot suggest anyclear synchronicity in glacial and climatic events. Furthermore, we introduce six additional locations where glacier moraines have been wave-washed or cut by postglacial raisedmarine shorelines, suggesting the landformswere deposited before or during high relative sea-level stands, thus exhibiting a similar LGEH age. Contrary to current understanding, our new evidence suggests that the LGEH glaciers were more dynamic, exhibited re-advances and extended well beyond the extensively studied LH glacial expansion. Given the widespread occurrence of the LGEH glacier deposits on Svalbard, we suggest that the culminationof theNeoglacial advancesduring theLittle IceAgedoesnotmark themaximumextentofmostSvalbard glaciers since deglaciation; it is just the most studied and most visible in the geological record. Wesley R. Farnsworth (WesleyF@unis.no), Olafur Ingolfsson, Michael Retelle and Lena M. Hakansson, University Centre in Svalbard (UNIS), Longyearbyen N-9171, Norway; Wesley R. Farnsworth, Lis Allaart and Anders Schomacker, UiT The Arctic University of Norway, TromsøN-9037, Norway; Olafur Ingolfsson, University of Iceland, Askja, Sturlugata 7, Reykjavık IS-101, Iceland; Michael Retelle, Bates College, Lewiston, ME 04240, USA; Anders Schomacker, Natural History Museum of Denmark, University of Copenhagen, Øster Voldgade 5-7, Copenhagen K DK-1350, Denmark; received 14thMarch 2018, accepted 3rd May 2018. Studying past changes in high latitude glaciers and climate allows us to better understand the role of the Arctic in the global climate systemand tomore effectively predict antecedent climate scenarios (McKay & Kauf- man 2014). Climatic shifts are occurring with an ampli- fiedeffect in theArcticatpresent andalsodid so following the last deglaciation (Hald et al. 2007; Hartmann et al. 2013). Arctic glaciers can be valuable indicators for past climate given their sensitivity to summer temperatures and winter precipitation (Huss & Hock 2015). Due to Svalbard’s high Arctic location between the northern extent of theNorthAtlanticDrift and the seasonal sea ice, the region is an ideal location to study the European Arctic’s sensitivity toglobal climate change (Fig. 1;Rogers et al. 2005). Basedonevidencefromcirqueglaciersonthewestcoast ofSpitsbergen,Svalbard,glaciersduring theLittleIceAge (LIA),culminating c. 1920CE,werebelievedtobeat their greatest extent since before the Younger Dryas (YD; Svendsen & Mangerud 1997; Mangerud & Landvik 2007). Studies fromwestern Spitsbergen suggest minimal glacier activity during the Lateglacial and earlyHolocene (LGEH), while the Neoglacial (c. 4.5 ka BP) cooling, which culminated during theLIA, has been interpreted to represent the Holocene glacial maximum in most loca- tions (Werner1993;Svendsen&Mangerud1997;Reusche et al.2014;Milleret al.2017). It isassumed that if glaciers had extended further than their late Holocene (LH) maximum, itwouldhavebeenduring theYD(Svendsen& Mangerud 1992). In a review of the deglaciation of Svalbard, Hormes et al. (2013) suggested that one of the key research questions still remaining is whether there is any evidence of YD climatic cooling seen in Svalbard. Until now, few pre-LH moraines have been identified and only one has been dated, where two moraine ridge boulders provide ambiguous cosmogenic exposure ages (12.1 and 9.2 kaBP;Henriksen et al. 2014).Additionally, a couple of earlyHolocene glacier re-advances (following deglaciation) have been inferred from stratigraphical data of geological sections (Lønne2005)or lake sediments (van der Bilt et al. 2015). With few exceptions, current evidence (and lack thereof) suggests that since the deglaciation of Svalbard, the most extensive glacier positions occurred in the LH and inmost cases roughly 100 years ago (Werner 1993; Humlum et al. 2005; Mangerud & Landvik 2007). The majority of Holocene glacier studies from Svalbard have focused on LH landforms, i.e. Neoglacial and Little Ice Age glacial deposits, which occur prominently andwidespread throughout the landscape (Werner 1988, 1993; Reusche et al. 2014; Philipps et al. 2017). However, as a result of ‘erosional censoring’ the preservation of glacial deposits in Svalbard is strongly biased towards the youngest deposits (Kirkbride &Winkler 2012; Landvik et al. 2014). Wepresent evidenceof a re-advanceofGullfaksebreen intoDeGeerbukta in northeast Spitsbergen constrained in time to the LGEH transition (Fig. 1). We summarize the glacial and palaeoclimatic conditions on Svalbard during the transition from glacial to interglacial mode between 14–9 ka BP (Landvik et al. 2014) and compare DOI 10.1111/bor.12326 © 2018 Collegium Boreas. Published by John Wiley & Sons Ltd bs_bs_banner the Gullfaksebreen re-advance to the timing of previ- ously published studies implying other LGEH glacier advances. We further up-scale this study by highlighting sixadditional sitesonSvalbardwhereglacierdeposits are cut by relatively high raised shorelines suggesting a similar LGEHage to the Faksebreen advance (Fig. 1A). Material and methods De Geerbukta fieldwork and sampling was conducted in August 2015, with focus on geomorphological and sedi- mentological mapping. Following field sampling, shell and whalebone subsamples were cleaned, weighed, pho- tographed, identified and subsequently sent for dating at theAngstr€omLaboratory at UppsalaUniversity.We pre- sent four new radiocarbon ages of shells and whalebone from De Geerbukta (Table 1). Additional LGEH glacier marginal deposits (mor- aines) have been identified on the open access website TopoSvalbard, provided by the Norwegian Polar Insti- tute. Access to modern (2008–12) high-resolution aerial imagery, digital elevationmodels and1936oblique aerial photographs provided the opportunity to thoroughly inspect the unglaciated landscape from numerous per- spectives (Fig. S1). Cross-cutting relationships between raised marine shorelines and glacier moraines indicate that deposits left at previous glacial margins occurred during a high relative sea-level stand (Ingolfsson & Landvik 2013) while maximum constraining ages are inferred from the deglaciation ages of the fjordor region. The principle of parsimony thus suggests the moraines were deposited following the deglaciation and prior to significant drop in relative sea level; thus, the age of the moraines roughly straddle the interval of time corre- sponding to the LGEH. Modern glacier lengths were extracted from Svalbardkartet (K€onig et al. 2014) while palaeo-glacier lengths were measured in TopoSvalbard (Table S1). Each 14C age of a marine organism presented in this study has been corrected and calibrated with IntCal13 (Reimer et al. 2013) using the Calib Rev. 7.0.4 program. A marine reservoir age has been implemented by select- ing ‘MARINE13’ and inputting a ‘Delta R’ of 7030 (Table 1; Mangerud & Svendsen 2018). All dates (of marineorganisms) presented inprevious studies havebeen re-(corrected and) calibrated accordingly (Table S2). The ages presented in text and figures are calibrated median ages in kilo-years before present (cal. ka BP). Results The Fakse moraine, De Geerbukta Gullfaksebreen is an outlet glacier of the Asgardfonna icecap located in NE Spitsbergen, extending 20 km and draining into De Geerbukta (Fig. 1A–C). Substantial ice-cored LH moraines protrude into Faksedalen from Gullfaksebreen and other outlet glaciers from Asgard- Fig. 1. A. Overview map of northern Spitsbergen. Symbols indicate sites with Lateglacial–early Holocene moraine ridges (see legend; referenced studies marked by grey triangles while Lateglacial–early Holocene glacier deposits marked by black dots, data presented in Tables S1, S2. B. Inset map of the North Atlantic, Svalbard marked by red box. C. Topographic map of Faksedalen with moraine deposits shaded grey. The entire Fakse moraine has been wave-washed indicated by dark lines while late Holocene deposits have not. Red circles and dotted line indicate distance measurements. White star marks the study site with proglacial glaciotectonized marine sediments presented in Fig. 2. Base maps in (A) and (C) modified from TopoSvalbard, (B) modified from IBCAO. 2 Wesley R. Farnsworth et al. BOREAS fonna (Fig. 1C). A 3.5-km-long arcuate moraine system (‘Fakse moraine’) with its highest point 52 m a.s.l. is located outside of the LH moraines at the mouth of Faksedalen (Fig. 2A–C; ArcticDEM). The Fakse mor- aine system is cut by theFaksemeltwater stream.Bedrock protrudes through the moraine material along the north- eastern part of the ridge. Glaciotectonized marine sedi- ments are exposed in a geological section in the distal flank,outboardof theFaksemoraine (Figs 2D–F,3).The glaciotectonized marine sediments were proximal to the ice margin but the sample site was not over-ridden (Figs 1C, 2D–F, 3). Following the retreat of Gullfakse- breen, the entire surface of themorainewaswave-washed, depositing shells and whalebones in the beach sediments (Fig. 2G–I; Table 1). The younger bivalve found in living position in the deformed soft marine (silts and sands) sediments on the distal side of the Fakse moraine dates to 11.9 cal. ka BP (11 930 cal. a BP), effectively constraining a maximum age for thedepositionof theFaksemoraine (Fig. 2G).We suggest these marine sediments were deposited in a shallow fjord environment (45–55 m water depth) with glaciers in the catchment but not directly proximal to a tidewater ice margin. The bivalve shells from the glacio- tectonized marine sediments pre-date the moraine and suggest thatGullfaksebreen advanced intoDeGeerbukta sometime after 11.9 cal. ka BP (Table 1). Shell fragments at 45 ma.s.l. and awhalebone at 28 ma.s.l. sampled from beach sediments overlaying the Fakse moraine give minimum constraining ages for the deposit, dating to 10.6 and9.6 cal. kaBP, respectively (10 550and9590 cal. a BP; Fig. 2H, I). Although the mollusc age dates that specific shoreline, the Fakse moraine, which has been entirely wave-washed, suggests the glacier had not only deposited the moraine, but retreated from its maximum position during the formation of the shoreline. Because the entire Fakse moraine has been inundated during a high relative sea level (above 52 m a.s.l.), we assume the glacier advanced into the fjord during or shortly after the maximum postglacial relative sea-level stand, sometime between 11.9–10.6 cal. ka BP (Fig. 3). The older bivalve found in living position in the glaciotectonized marine sediments has a median age of 12.2 cal. ka BP (12 190 cal. a BP; Fig. 3, Table 1). The shells sampled from the deformed sediments are close in age, making it unclear how long the outer extent of De Geerbukta was ice-free before the Gullfaksebreen re- advance. Interestingly, the median age of the older bivalve extends the minimum age of the deglaciation of the Hinlopen Strait from roughly 11.3 to 12.2 ka BP (Hogan et al. 2017; Flink et al. 2017). Extended Lateglacial–early Holocene glaciers Based on the interpretation of surficial deposits visible in aerial imagery and digital elevation models of TopoSval- bard we introduce six locations where LGEH glaciers re-Ta bl e 1. D e G ee rb uk ta ra di oc ar bo n da te s. D at es co rr ec te d fo r m ar in e re se rv oi r ag e w it h D R = 70 3 0 (M an ge ru d & Sv en ds en 20 18 )a nd ca lib ra te d w it h M ar in e1 3 on C al ib R ev 7. 0. 4. L ab .I D L oc at io n D at e Sa m pl e D ep os it E le va ti on (m a. s. l.) L at ./L on g. (° ) 14 C ag e (a B P ) d1 3 C co rr .( & ) C al .a ge (a B P )M ed ia n C al .a ge (a B P )1 r U a- 52 52 1 D e G ee rb uk ta 12 .8 .2 01 5 M . tr un ca ta fr ag . B ea ch sa nd an d gr av el 45 79 .6 01 40 0 N ,1 7. 65 50 00 E ic e 97 25 4 2 1 10 55 3 10 48 4– 10 62 2 U a- 52 52 0 D e G ee rb uk ta 12 .8 .2 01 5 W ha le bo ne B ea ch sa nd an d gr av el 28 79 .6 13 20 0 N ,1 7. 68 60 00 E 90 20 3 5 1 8. 5 95 91 95 30 –9 65 2 U a- 52 52 2 D e G ee rb uk ta 12 .8 .2 01 5 M . ca lc ar ea pa ir ed sh el l D ef or m ed sa nd 8 79 .6 19 48 0 N ,1 7. 67 21 50 E 10 68 6 43 0 .1 11 93 3 11 82 6– 12 03 9 U a- 52 52 3 D e G ee rb uk ta 12 .8 .2 01 5 M . ca lc ar ea pa ir ed sh el l D ef or m ed sa nd 7 79 .6 19 48 0 N ,1 7. 67 21 50 E 10 82 7 44 1 .5 12 19 3 12 05 9– 12 32 6 U a- 52 51 8 P al an de rb uk ta 10 .8 .2 01 5 H . ar ct ic a fr ag . B ea ch sa nd an d gr av el 53 79 .5 60 44 0 N ,2 0. 62 42 50 E 98 94 4 0 1. 4 10 74 2 10 65 5– 10 82 9 BOREAS Svalbard glaciers re-advanced during the Pleistocene–Holocene transition 3 Fig. 2. A–C.OverviewoftheFaksemorainewithalluvialsystemandDeGeerbuktaontheglacialproximalanddistalsiderespectively.D–F.Glaciotectonized marine sediments located on the distal side of the Fakse moraine. Deformed structures outlined for clarity. Location of the proglacial deformed sediments markedbywhite star inFig. 1.G–I. Sampledpaired shells in glaciotectonizedmarine sediments aswell as shell fragments andwhalebone inbeach sediments overlaying the Faksemoraine, respectively. Corrected/calibrated ages and elevations indicated. 4 Wesley R. Farnsworth et al. BOREAS advanced or exhibited stillstands substantial enough to leave remnant moraine ridges (Figs 1A, 4, S1; Table S1). Similar to the Faksemoraine inDeGeerbukta, identified ridge deposits are interpreted to represent ice-marginal deposits often in the form of terminal (or frontal-lateral) moraines. In all cases, themoraine ridges are located at or around the marine limit. Glacial deposits have been partially or entirely wave-washed allowing for the poten- tial to constrain the event with detailed sea-level curves (Forman 1989; Fig. S1). Palaeo-glacier lengths for these six glaciers aswell asDe Geerbukta extended between roughly 2–8 km beyond their LH margins (Figs 1C, 5). On average, the LGEH glacial deposits that we introduce are ~4 km outboard of the LH moraines (Figs 1C, 5; Table S1). The sites with glacial deposits of suggested LGEH age are found across Spitsbergen and correspond to glaciers of all sizes with modern lengths ~1–20 km (Fig. 5). The two larger glaciers,Gullfaksebreen andTabreen, drain from thehigh plateau icecap, Asgardfonna. Glaciers Heftybreen and Albertbreen correspond to small cirques, while Flower- breen and Richterbreen today are small cirque glaciers; but during the LGEH both were most likely valley glaciers. The glacier that filled Lyckholmdalen during the LGEH probably was an outlet glacier from an icecap, whichmaycorrespond to themodernFriggkapa(Fig. S1). In some cases, glacial deposits correlate to empty catch- ments, like Heftybreen, where at present the cirque is almost entirely ice-free (Fig. S1). Although an unknown amount of ice, remnant from the degrading Svalbard Barents Sea Ice Sheet (SBIS) probably still existed, all sites with LGEH moraines suggest transitional or local flow styles confining within topography and overprinting maximum ice-flow directions (Landvik et al. 2014). Discussion Lateglacial–early Holocene conditions and glacier deposits Between14–12 cal.kaBP,extensiveregionsofthemarine- basedsectorsof theSBIScollapsed,resulting inaretreatof thenorthwesternmargin from the shelf topositions inside the modern coastline of Svalbard (Mangerud et al. 1992; Hormes et al. 2013). Deglaciation was time transgressive and initially occurred in major troughs and fjords and subsequently in the inter-trough and terrestrial realms (Ingolfsson & Landvik 2013). The Holocene Thermal Optimum (HTO) is characterized by the migration and presence of warm water molluscs and foraminifera to waters around Svalbard between 11.2–5.5 cal. ka BP (Fig. 5; Salvigsen 2002; Hald et al. 2004). Proxy records fromlakesedimentscollected innorthwesternSpitsbergen suggest warm, moist conditions as early as 12.8 cal. ka BP (Balascio et al. 2018; van der Bilt et al. 2018) while a recent review of thermophilous molluscs suggests early warmconditionswithpeakwarming (6 °Cgreater thanat present) from10.2–9.2 cal. kaBP(Mangerud&Svendsen 2018). Marine terminating glaciers responded to the warming by retreating back to terrestrial margins while some cirque glaciers disappeared completely (Snyder et al. 2000; Forwick & Vorren 2009). A study highlighting transitional deposits on Spitsber- gen suggests that there are numerous terrestrial and subaqueous glacial deposits that could be related to glacier advances during this period, traditionally associ- ated with the deglaciation (Landvik et al. 2014). The LGEH glacier advances have been noted in some strati- graphical studies from Spitsbergen, but have been the Fig. 3. Schematic cross-section of theFaksemoraine. Corrected/calibrated radiocarbonages constrain the glacier advance in time. BOREAS Svalbard glaciers re-advanced during the Pleistocene–Holocene transition 5 focusoffewerstudiescompared totheLHglacierdeposits (Table S1). Studies in central Spitsbergen describe poorly constrained LGEH glacial deposits in Bromelldalen (Fig. 1A; site 18) and Gangdalen (Fig. 1A; site 21) but focus on regional sea levels, marine sediments and the deglaciation (Landvik&Salvigsen 1985;Mangerud et al. 1992). Additionally, shoreline studies from north central Spitsbergen partly constrain LGEH glacial deposits for several catchments (Fig. 1A; sites 1–8; Br€uckner et al. 2002; Eitel et al. 2002). More recently, stratigraphical sections near Longyearbyen (Fig. 1A; site 17) have been interpreted to represent a valley glacier advance between 10.6–10.4 cal. ka BP (Fig. 6; Lønne 2005). The deposits are located ~7 km outside of the present ice margin, and are at the regional Holocene marine limit (63 m a.s.l.). Lønne (2005) links this advance to a regional climatic Fig. 4. High-resolution aerial imageryof six selected siteswhere glacier deposits (marked bywhite dotted lines) havebeen remotely constrained to Lateglacial–early Holocene age. Glacier deposits labelled A–F corresponding to the Tabreen, Albertbreen, Lyckholmdalen, Flowerbreen, Heftybreen and Richterbreen glaciers/valleys, respectively. Glacier deposits presented in (A) and (C) correspond to outlet glaciers from plateau icecaps and (B) and (E) from cirque glaciers, while (D) and (F) highlight glacier deposits from valley glaciers. Imagery from TopoSvalbard. For topographic data and larger images see Figs S1, S2. 6 Wesley R. Farnsworth et al. BOREAS event, comparing other local advances in the Isfjorden area such as an advance of Aldegondabreen (SW Isfjor- den) subsequent to 10.9 cal. ka BP and an advance of Esmarkbreen (N Isfjorden) subsequent to 10.3 cal. kaBP (Salvigsen et al. 1990; Lønne 2005). New data from St. Jonsfjorden suggest a tributary glacier advanced between 12.8–10.1 cal. ka BP (Fig. 5; site 15; Farnsworth et al. 2017).Theoldestmaximumconstraining agesof reported glacier deposits reflecting postglacial re-advances corre- spond to two studies on north and south Prince Karls Forland with dates of 14 and 13.1 ka BP, respectively (Fig. 1A, sites 13, 14; Andersson et al. 1999, 2000). The youngest date reported as a maximum constraining age is fromEdgeøya, located in the east ofSvalbardwhere glacial sediments (withawhalerib)outboardof theAlbrechtbreen Neoglacialdeposits date to9.2 cal. kaBP(Fig. 1A; site 23; Ronnert & Landvik 1993). Recent mapping in the Van Mijenfjorden identifies glacial deposits connected to a delta remnant from a high relative sea level, also suggested to be earlyHolocene in age (site 20; Larsen et al. in press). Van der Bilt et al. (2015) suggest a Holocene (glacial) maximum between 9.6–9.5 cal. ka BP, based on a lake sediment core fromNWSpitsbergenand suggest links toa meltwater pulse from the melting Laurentide Ice Sheet (LIS) via the Hudson Strait influencing regional climate (Fig. 1A; site 10). In contrast with the suggested early Holocene glacial maximum interpreted from the lake Hajeren sedimentary record, several studies suggest this period in time coincides with peak warming in Svalbard fjords and regional waters (Hald et al. 2004;Mangerud& Svendsen 2018). Although van der Bilt et al. (2015) assumethat thepeak inminerogenic-rich sedimentsdating from 9.6 to 9.5 ka BP in the lacustrine chronology reflects glacier advances from two snowpatches driven by a large- scale connection to the melting LIS, other deposits in the catchment may also influence the Hajeren lake record. Interestingly, a study from Kløsa, the lake just one kilometer to the southwest of Hajeren suggests the outer Karlbreen moraine pre-dates the HTO. Interestingly, a study from Kløsa, the lake just one kilometer to the southwest of Hajeren suggests the outer Karlbreen morainepre-dates theHTO.Given themappedcatchment boundary for lake Hajeren was drawn over the Karlbreen moraine in the southern end of the drainage (van der Bilt et al. 2015, 2016); it seems a simpler solutions for the peak in minerogenic sediments from Hajeren are sourced from thedegradationof theKarlbreenmorainedrivenbywarm regional conditions (Hald et al. 2004; Mangerud & Svendsen 2018). Although the interpretation of the chronology from Hajeren may not be sound, this would still suggest the Karlbreen moraine pre-dates the early Holoceneglaciermaximumand formedsome-timeduring the LGEH (van der Bilt et al. 2015; Røthe et al. 2015). In summary, most of these deposits are poorly constrained in age and seem to have formed between 14.0–9.0 cal. ka BP without clear synchronicity or prominent terminal moraines (Fig. 6). Thus, if they are related, it is not directly through one specific climatic event, but indirectly owing to environmental or external factors. Additionally, marine palaeoclimate records suggest that regional conditionsweremuchwarmer than at present (Hald et al. 2004; Mangerud & Svendsen 2018). Glacier re-advances during deglaciation Our current understanding of the regional deglaciation and Holocene history of Svalbard has only involved terminal moraines of re-advancing glaciers during the LH (Mangerud & Landvik 2007; Hormes et al. 2013). We revise this reconstruction by presenting the first well- constrained end moraine complex formed during the Lateglacial–early Holocene in De Geerbukta, NE Sval- bard.We propose that the Fakse moraine was deposited in a fjord during a re-advance between 11.9–10.6 cal. ka BP. At that time, the palaeo-Gullfaksebreen extended over 5.5 km downvalley from the modern ice margin (4.4 km past the LH), marking the outermost glacier extent since the deglaciation. The age constraint of theFaksemoraine and respective re-advance straddling the transition from the Pleistocene toHolocene equates to either the largestHolocene glacier extent occurring during the early Holocene (not during the LIA), or stands as the best candidate for a glacier re- advance on Svalbard occurring during the YD chrono- zone. Both scenarios improve our understanding of the Fig. 5. Distances of Lateglacial–early Holocene glacier margins com- pared in length to the lateHolocenemaximumextent and themodern ice front (K€onig et al. 2014). Red dots represent Faksedalen (introduced in Fig. 1C).Greydotsrepresent the sixglacial systems introduced inFig. 4. In all casesLateglacial–earlyHolocene ice-marginal deposits are located several km outboard of the late Holocene glacier margins suggesting larger glaciers systems during the Lateglacial–early Holocene. For data see Table S1. BOREAS Svalbard glaciers re-advanced during the Pleistocene–Holocene transition 7 recent history of glaciers and icecaps on Svalbard. Furthermore, themaximumconstraining ages of bivalves fromtheglaciotectonized shallowmarine sediments inDe Geerbukta are the oldest radiocarbon ages inside of the Hinlopen Strait, pushing back the age of deglaciation a millennium in time (Table 1; Hogan et al. 2017; Flink et al. 2017). Additionally, in this study we summarize the synchro- nicity of our well-constrained end moraine to an addi- tional15othersitesmentionedinthe literature.Wefurther upscale these findings by highlighting six other locations where geomorphologically similar deposits exist terrestri- ally on Svalbard. These LGEH glacier deposits corre- spond toglaciers of all sizes, and the deposits often extend several kilometres beyondLHmaxima. Three of these six sites, which were identified remotely (in addition to De Geerbukta), have been ground-truthed, and as of yet, all lack the age constraint of the LGEH re-advance of Gullfaksebreen. We are confident in the general mor- phology of the moraine ridges and their location at the mouthof catchments, aswell as their relationship to older postglacial sea levels. Similar to the referenced study sites, the glacierdeposits commonly occur proximal to the postglacial high relative sea-level stands. More research is needed to determine whether the marine tidewater setting played a dynamic role or wasmerely a result of the environment at the time. Studies of the sea-floor morphology would be comple- mentary and potentially indicate whether these deposits are strictly linkedtotheLGEHrelative sea level.However, unlike the terrestrial LGEH glacier deposits that can be indirectly constrained in age remotely, the submarine glacier deposits do not exhibit cross-cutting morpholog- ical evidence of a high relative sea level (constituting a minimum constraining age of deposition). Recent work from eastern Spitsbergen suggests glaciers may have extended further than the traditional Neoglacial–LIA maxima during the early Holocene, but the moraines are yet tobewell constrainedintime(Flink&Noormets2017; Flink et al. 2018). Fig. 6. Diagram summarizing Lateglacial–earlyHolocene climate and glacier deposits in Svalbard. d18O fromNGRIP plotted in blue (Andersen et al. 2004). Summer insolation at 79°N indicated by red curve and sea-ice extent indicated by horizontal solid and fragmented blue lines based on proxy IP25 (M€uller&Stein 2014). Black barsmark constrained glacier events, dashed lines indicate lacustrine studies anddotted lines indicate age range of dated glacial deposits. Partially constrained eventsmarked byT-forms.Red/light blue bars indicatewarm/cool conditions based onwarm water species and proxies. Thick red line indicates maximum fjord temperatures at least 6 °C greater then present (Mangerud & Svendsen 2018). Shaded vertical bar marks the Younger Dryas period. See Table S2 for raw data and references. 8 Wesley R. Farnsworth et al. BOREAS Glacier re-advances of LGEH age are not unique to Svalbard andhavebeen identifiedat otherNorthAtlantic ice-sheetmarginsduringthe lastdeglaciation.Someoutlet glacier re-advances have been identified in Iceland shortly following thecollapseof themarine-terminatingIcelandic Ice Sheet between 13.4–11.7 cal. ka BP (Norðdahl & Ingolfsson 2015; Patton et al. 2017; Sigfusdottir et al. 2018). Local glaciers in North Greenland re-advanced in response to theYDcooling as interpreted from theGRIP ice-cores (Larsen et al. 2016),whereasmost otherglaciers in Greenland did not advance during the YD period (Denton et al. 2005; Funder et al. 2011). An extremely dynamic deglaciation is visible in offshore data from northern Scandinavia and the Barents Sea, where the interplay between rising eustatic sea level and grounding- line conditions influenced retreat, re-advances and stag- nation (Newton & Huuse 2017). Furthermore, a non- linear deglaciation of the NW margin of the Laurentide Ice Sheet is suggested to have been interrupted by numerous re-advances and variable retreat rates (Stokes et al. 2009). It appears that as ice flow transitions from maximum to local flow style (Landvik et al. 2014) at the margins of ice sheets during deglaciation, glacier beha- viour is increasingly being characterized by not just retreat, but by unsustainable (dynamic) re-advances, stagnation and episodic retreat. It is possible this may reflect the style of deglaciation and provide an analogue for how large bodies of ice rapidly shed mass, thus furthering our understanding of deglaciation. As most glacier advances are mass balance controlled, driven by increased winter precipitation and/or lower summer temperatures, deposits from advancing glaciers constrained in age are associated with periods of cooler summer temperatures and/or abundant snowfall (Solom- ina et al. 2015). But, is this always the case? With our current understanding of the palaeoclimate of Svalbard one cannot test whether palaeo-glacier advances are forced by winter precipitation, summer temperatures, or internal glacier dynamics. Ifmassbalance is the keydriver of these glacier advances, a rough regional synchronicity in glacier fluctuationswould be expected (which is not yet visible). Thus, it is crucial to constrain more than a single glacier to effectively infer past climatic conditions. As it now appears, the time transgressive deglaciation of the fjord-systems may play an indirect role in the timing of the glacier re-advances. Thus, a re-advance located near the mouth of the fjord (i.e. Heftybreen, Site E)may be older than the age of themoraine located in an inner tributary fjord (Lyckholmdalen, Site C). And in thismodel, amid-fjord tributary likeBolterdalen (Lønne 2005)mayhavere-advancedat somepoint inbetween the timing of the other two glaciers’ re-advances. OnSvalbard, theLGEHdeposits correspond toglaciers of all sizes. Additionally, numerous glacier deposits relate to currently empty cirques and valleys suggesting a connection to the deglaciation dynamics, and the subse- quent Holocene climate not being favourable enough to reinvigorate the glaciers following their unsustainable advances (Figs 4, S1). Given the high frequency of surge- type behaviour seen on Svalbard (Sevestre & Benn 2015; Farnsworth et al. 2016), the simplest reasoning would suggest that during the LGEH glaciers also may have exhibited surge-type behaviour. Furthermore, several stratigraphical studies from western Spitsbergen have highlighted glacial re-advances from earlier in the Pleis- tocene during periods assumed to be warm and charac- terized by deglaciation (Fig. 1A; sites 13, 14 and 22; Landvik et al. 1992; Andersson et al. 1999, 2000). At present, the lack of synchronicity and the unknown rate of glacier advance/retreat corresponding to most LGEH glacier moraines make it challenging to discern whether the glacial oscillations that left the deposits were controlled by climate, internal glacier dynamics (surges), or a combination of these factors. We need to better understand: (i) past winter precipi- tation and summer temperatures, (ii) rapid glacio- isostatic uplift rates and their potential influence on mass balance, (iii) sea ice–glacier interactions, and (iv) the dynamics of the break-up of an ice sheet, to identify casual linkages to climate or internal dynamics. The transition from the Lateglacial to early Holocene resulted in glacier re-advances, stillstands and complex behaviour at many ice-sheet margins during the past deglaciation. It is reasonable to expect the reorganizing of glacier hydrological systems, shifting thermal regimes and fluctuating glacier surface profiles during the LGEH on Svalbard (Ingolfsson & Landvik 2013; Landvik et al. 2013, 2014). Considering the increasing extent of open water, warm fjords, relatively high sea levels and summer insolation characteristic of the deglaciation, it is not surprising that glaciers exhibited complex behaviour. Conclusions Glacier deposits of LGEH age are present across Sval- bard, correspond to glaciers of all sizes and suggest a periodofmuchmoreactiveglaciodynamicbehaviour than previously understood. These LGEH moraines extend significantly further than LH deposits that were previ- ously believed to reflect the most extensive glacier extent on Svalbard since the last deglaciation. The LGEH deposits described here provide insight into deglaciation dynamics and the behaviour of marine-terminating ice sheets in the changing climate of the Pleistocene to the Holocene transition. These deposits reflect hiccups in the overall deglaciation on Svalbard. Although it is unknown whether theseglacierdeposits relate toclimaticordynamic forcing, they do suggest the style of the deglaciation of ice sheets may be more complex than our current under- standing. Thus, deglaciation may not reflect a simple and continuous retreat, but rather be characterized by re- advance, stagnation and subsequent recession (Newton& Huuse 2017). BOREAS Svalbard glaciers re-advanced during the Pleistocene–Holocene transition 9 Acknowledgements. – The authors acknowledge Sveinn Brynjolfsson and Sara Mollie Cohen for assistance in the field. Fieldwork and radiocarbondateswerefundedbytheCarlsbergFoundation(CF14-0756 to AS) and UNIS, respectively. 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Quaternary Science Reviews 111, 9–34. Stokes, C. R., Clark, C. D. & Storrar, R. 2009: Major changes in ice streamdynamics during deglaciation of the north-westernmargin of the Laurentide Ice Sheet.Quaternary Science Reviews 28, 721–738. Svendsen, J. I. & Mangerud, J. 1992: Paleoclimatic inferences from glacial fluctuationsonSvalbardduring the last 20 000 years.Climate Dynamics 6, 213–220. Svendsen, J. I. & Mangerud, J. 1997: Holocene glacial and climatic variations on Spitsbergen, Svalbard. The Holocene 7, 45–57. Werner, A. 1988:Holocene glaciation and climatic change, Spitsbergen, Svalbard. Ph.D. thesis, University of Colorado, 296 pp. Werner, A. 1993: Holocene moraine chronology, Svalbard: licheno- metric evidence for multiple neoglacial advances in the Arctic. The Holocene 3, 128–137. Supporting Information Additional Supporting Informationmay be found in the online version of this article at http://www.boreas.dk. Fig. S1. TopoSvalbard maps of Lateglacial–early Holo- cene sites. Fig. S2. Enlarged imageryofLateglacial–earlyHolocene sites. Table S1.Glacier length data. Table S2. Lateglacial–early Holocene studies. BOREAS Svalbard glaciers re-advanced during the Pleistocene–Holocene transition 11 Chapter II Lateglacial and Holocene glacier activity in the Van Mijenfjorden area, western Svalbard Larsen, E., Lyså, A., Rubensdotter, L., Farnsworth, W.R., Jensen, M., Nadeau, M., Ottesen, D. (2018): Lateglacial and Holocene glacier activity in the Van Mijenfjorden area, western Svalbard. arktos. DOI: 10.1007/s41063-018-0042-2. Oblique aerial image of the mouth of Gustavdalen, northern Van Mijenfjorden, Spitsbergen. Chapter III Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard Farnsworth, W.R., Ingólfsson, Ó., Noormets, R., Allaart, L., Alexanderson, H., Henriksen, M., Schomacker, A. (2017): Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard. Boreas 46, 585-603. DOI: 10.1111/bor.12269. Modiolus modiolus sampled at Piriepynten, inner St. Jonsfjorden, Spitsbergen, July 2015. Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard WESLEY R. FARNSWORTH , OLAFUR INGOLFSSON, RIKO NOORMETS, LIS ALLAART, HELENA ALEXANDERSON, MONA HENRIKSEN AND ANDERS SCHOMACKER Farnsworth, W. R., Ingolfsson, O., Noormets, R., Allaart, L., Alexanderson, H., Henriksen, M. & Schomacker, A. 2017 (July): DynamicHolocene glacial historyof St. Jonsfjorden, Svalbard.Boreas, Vol. 46, pp. 585–603. https://doi. org/10.1111/bor.12269. ISSN 0300-9483. Evidenceof a dynamicHoloceneglacial history is preserved in the terrestrial andmarine archives of St. Jonsfjorden, a small fjord-systemon thewest coastof Spitsbergen, Svalbard.High-resolution, remotely sensed imagery frommarine andterrestrialenvironmentswasusedtoconstructgeomorphologicalmaps thathighlightan intricateglacialhistoryof the entire fjord-system. The geomorphology and stratigraphy indicate an early Holocene local glacier advance constrained to the Lateglacial–earlyHolocene transition. Identification and 14C dating of the thermophilous bivalve molluscModiolus modiolus to 10.00.12 cal. ka BP suggest a rapid northward migration of the species shortly after deglaciation. Further evidence enhances the understanding of the onset and subsequent climax of the Neoglacial- Little IceAge in innerSt. Jonsfjorden.Thepresent-day terminusofOsbornebreen, thedominatingglacier systeminSt. Jonsfjorden, is locatedover8.5 kmup-fjord from itsNeoglacialmaximumextent.Cross-cutting relationships suggest subsequent advances of all the smaller glaciers in the area following the break-up of Osbornebreen. Glacial deposits, landforms and their cross-cutting relationships observed in both terrestrial andmarine settings imply a complex and highly dynamic environment through the later part of the Holocene. Wesley R. Farnsworth (WesleyF@unis.no), Olafur Ingolfsson and Riko Noormets, University Centre in Svalbard (UNIS), Longyearbyen N-9171, Norway;Wesley R. Farnsworth, Lis Allaart and Anders Schomacker, UiT The Arctic UniversityofNorway,TromsøN-9037,Norway; Olafur Ingolfsson,Universityof Iceland,Askja,Sturlugata7,Reykjavık IS-101, Iceland; Helena Alexanderson, Lund University, S€olvegatan 12, Lund 223 62, Sweden; Mona Henriksen, Norwegian University of Life Sciences, P.O. Box 5003, Ås N-1432, Norway; received 3rd February 2017, accepted 19thMay 2017. The climate during the Holocene (the last 11.7 ka) has traditionally been regarded as relatively stable, com- pared to the preceding Late Pleistocene (Dansgaard et al. 1993; Steffensen et al. 2008; Rockstr€om et al. 2009), but this paradigm of a relatively uneventful Holocene is being increasingly challenged (Bond et al. 2001; Mayewski et al. 2004; Wanner et al. 2011). We are beginning to see that shifts in climate occur more rapidly than previously understood, particularly in Arctic regions at present and during the early Holocene due to Arctic amplification (Bond et al. 2001; Hald et al. 2007; Wanner et al. 2011; Masson- Delmotte et al. 2013; Briner 2016). By studying the rate and magnitude of past changes in climate we are able to better understand current climate dynamics, as well as more effectively predict antecedent climate scenarios (McKay & Kaufman 2014). Arctic glaciers can be valuable indicators for past climate variations given their sensitivity to winter precipitation and summertime temperatures (Oerlemans 2005). Not all glaciers respond uniformly to shifts in climate (Røthe et al. 2015), and in some cases ice-front fluctuations can be influenced by internal dynamics (surge beha- viour), making the interpretation of their climatic relationships challenging (Yde & Paasche 2010; Farns- worth et al. 2016). Additionally, marine terminating glaciers are subject to more complex glaciodynamic conditions compared to terrestrial systems, as ice-front fluctuations are influenced by oceanic temperatures, currents, sea ice and relative sea level as well as air temperature and precipitation (Forwick & Vorren 2009; Nick et al. 2010). The position of Svalbard between the northern extent of the North Atlantic Drift and the seasonal sea ice makes it an ideal setting to study the climate sensitivityof the Arctic (Rogers et al. 2005). In addition, Svalbard’s proximity to the former margin of the Svalbard-Barents Sea Ice Sheet (SBIS) has made it a valuable locality to study the configuration, behaviour and dynamics of the ice sheet (Landvik et al. 1998, 2005, 2014; Mangerud et al. 1998; Hormes et al. 2013; Ingolfsson & Landvik 2013; Gjermundsen et al. 2015). TheHolocene develop- ment of climate and glaciers in Svalbard is not well understood, partly due to the lack of case studies and, consequently, the scarceness of well-constrained (pre- Little Ice Age) chronologies for glacial events. The few detailed studies have predominantly focused on the west coast and have been somewhat biased towards younger events (Landvik et al. 2014), such as the Neoglacial and Little Ice Age (LIA) (e.g. Snyder et al. 2000; Humlum et al. 2005;Mangerud&Landvik 2007; Holmgren et al. 2010; Miller et al. 2017). As many Svalbard outlet glaciers terminate in fjords it is important to combine data from terrestrial and marine archives to holistically understand Holocene glacier oscillations and to make inferences about causal links in the climate-glacier system. The aim of this study is to reconstruct a dynamic and complex series of glacial events during the Holocene that are preserved in sedimentary and morphological DOI 10.1111/bor.12269 © 2017 Collegium Boreas. Published by John Wiley & Sons Ltd bs_bs_banner archives in a fjord system on the west coast of Spitsber- gen. We present this through detailed Quaternary geo- logicalmapsof terrestrial andmarineenvironments from inner St. Jonsfjorden. The Holocene history is chrono- logically constrained by 14C ages from stratigraphy, and cross-cutting relationships between glacial and marine landforms. Holocene history of Svalbard The Lateglacial–early Holocene on Svalbard is charac- terized by a rapid transition from aglacial to interglacial environment (Ingolfsson 2011; Hormes et al. 2013; Landvik et al. 2014). Between 15–12 ka BP extensive regions of the marine-based sectors of the SBIS col- lapsed,which resulted ina retreatof theNWmargin from the shelf area into the modern coastline of the Svalbard archipelago (Jessen et al. 2010; Hormes et al. 2013; Hughes et al. 2016). Deglaciation on Svalbardwas time transgressive andoccurred initially inmajor troughs and fjords and subsequently in the inter-trough and terres- trial realms (Hormes et al. 2013; Ingolfsson & Landvik 2013; Landvik et al. 2014). TheHolocene ThermalOptimum (HTO) on Svalbard is recorded first in the marine environment and subse- quently on land. The HTO is fingerprinted by the migration and presence of thermophilous molluscs and warm water foraminifera to waters around Svalbard between 11.2–5.5 ka BP (Salvigsen et al. 1992; Salvigsen 2002; Sarnthein et al. 2003; Hald et al. 2004; Blake 2006). This phase of warm stable conditions has also been interpreted from low sediment input and high productivity within lake basins between 8.5–5 ka BP (Birks 1991; Svendsen & Mangerud 1997; Holmgren et al. 2010; van der Bilt et al. 2015, 2017; Alsos et al. 2016). In response to thewarmer conditions, some of the marine terminating glaciers retreated onto land, while some terrestrial cirque glaciers may have melted away completely (Svendsen & Mangerud 1997; Snyder et al. 2000; Forwick et al. 2010; van der Bilt et al. 2015). Similar to the HTO, evidence from terrestrial and lacustrine environments related to Neoglacial climatic deterioration generally lags oceanic cooling. The decrease in abundance and number of species of ther- mophilous molluscs and foraminifera suggests that the marine cooling may have initiated as early as 8 ka BP (Salvigsen et al. 1992; Hald et al. 2004). Terrestrial and lacustrine studieshighlightglacier advances thatoccur in sporadic pulses between the period ~4.5–3 ka BP (Svendsen & Mangerud 1997; Reusche et al. 2014; van der Bilt et al. 2015; Røthe et al. 2015). The final pulse of the Neoglacial expansion of the glacial system is synchronous with the culmination of the LIA on Svalbard (~1920 CE), and is in general suggested to be the largest and most significant glacial phase of the Holocene (Werner 1993; Humlum et al. 2005). The culmination of the LIA is visible in historical 1936 Norwegian Polar Institute oblique aerial photographs that displaymost icemargins up against or just proximal to their LIAmoraine systems (Boulton et al. 1999; Lysa & Lønne 2001; Sletten et al. 2001; Paasche & Bakke 2010; Midgley & Tonkin 2017). Widespread retreat and thinning of Svalbard glaciers have occurred since the end of the LIA, and extensive areas of recently glaciated landscape have been revealed (Nuth et al. 2013). In some cases, tidewater glaciers have responded similarly to thewarmer conditions during the HTOand retreated backonto topographical heights and terrestrial environments (Forwick et al. 2010).Although the majority of glaciers have thinned and retreated in response to the 20th century climate, some glacier systems have exhibited rapid advances or speed-ups during this time, simulating the behaviour of a potential surge-type glacier (Sund et al. 2009; Lønne 2014; Dunse et al. 2015; Lovell et al. 2015a; Sevestre 2015; Farns- worth et al. 2016). These glaciers have subsequently thinned and retreated back at a considerably higher rate following their advances. Regional setting and climate Svalbard is an archipelago located along the dominant corridor of atmospheric moisture between the Atlantic and the Arctic Basin, spanning from 74° to 81°N (Vik- hamar-Schuler et al. 2016; Fig. 1A). The islands are positioned at the northern extent of the Gulf Stream (North Atlantic Drift) and the southern border of the Arctic sea-ice front (Rogers et al. 2005). Despite the archipelago’s northern location, western Spitsbergen experiences amild climate for its latitudewhere thewarm ocean current influences weather and sea ice (Førland et al. 1997, 2011). The climate is characterized by the interactions between the Icelandic Low and Siberian High pressure systems, and as a result, relatively high temperatures aredriven north overSvalbardby themain North Atlantic cyclone track (Hanssen-Bauer et al. 1990). Svalbard precipitation is closely coupled to the mode of the North Atlantic Oscillation (NAO; Dickson et al. 2000) and falls predominantly in solid form. St. Jonsfjorden is located in western Spitsbergen and opens into Forlandsundet, the sound that separates Spitsbergen from Prins Karls Forland (Fig. 1B, C). St. Jonsfjorden is 23 km long and ranges from 2 to 5 km in width (Fig. 1C). Nineteen glaciers drain into the fjord andat present three of themhavemarine termini,while in 1936atotaloffiveweremarineterminating.Thelargestof theglaciers foundat theheadof the fjord isOsbornebreen with an ice front that is >2 kmwide (Fig. 1C).Mountain peaks in the region range between 500–600 m a.s.l. in the outer fjord and up to 700–900 m a.s.l. in the inner fjord. The depth of the fjord basin is generally <100 m but exceeds130 min someareas in theouterpartof thebasin. The bedrock is diverse and ranges from Precambrian– Neoproterozoic marble, phyllite, quartzite and carbon- 586 Wesley R. Farnsworth et al. BOREAS ates, to Ordovician–Silurian sandstones and shales (Dallmann & Elvevold 2015). St. Jonsfjorden has been visited by whalers and sealers since the early 17th century, and the first scientific exploration of the area was carried out by the Isachsen Spitsbergen Expedition of 1909–10, where the bedrock geology was surveyed and the terminal position of Osbornebreen, themain tidewater terminus,was recorded (Hoel et al.1915–1917;Hoel 1929).Over thepast century, anumberof studies of raisedmarine sedimentsandglacial deposits in St. Jonsfjorden have highlighted past relative sea levels and the glacial history of the area (Feyling- Hanssen & Jørstad 1950; Dinely 1953; Forman 1989; Evans & Rea 2005). Two sets of shorelines have been identified, with an older setoccurringataround60 ma.s.l. close towhere the fjord opens into Forlandsundet and dating to the pre- Late Weichselian (Forman 1989). The younger set con- sistsofHolocene shorelines,where thehighestare located ~45 ma.s.l. inagreementwiththeLateWeichselian–early Holocene upper marine limit of the area (Forman 1989; Evans&Rea 2005).Minimumages of the deglaciation of the outer (fjord mouth) and inner (Piriepynten) parts of St. Jonsfjorden correspond to ~12.8 and 8.6 cal. ka BP, respectively (Forman 1989; Bunin 2015). Additionally, a Neoglacial maximum advance has been constrained by a shell-rich, coast-parallel, moraine with a maximum constraining age of 0.520.07 cal. BP (Evans & Rea 2005; Fig. 2). Methods and data acquisition Fieldworkandmollusc shell samplingwere conducted in July 2013, 2014, and 2015, with focus on geomorpho- Fig. 1. A. Map of North Atlantic with Svalbard boxed in red (base map modified from www.ngdc.noaa). B. Regional topographical map of Svalbard with St. Jonsfjorden boxed in red. Longyearbyen is marked by a star and major islands Prins Karls Forland (PKF), Spitsbergen and Nordaustlandet are identified. C. Topographical map of St. Jonsfjorden, with overview frames indicatingmapped regions presented in this study. Major tributaryglaciers aremarked. Topographicalmapsmodified fromTopoSvalbard (2017) provided by theNorwegian Polar Institute (2017). [Colour figure can be viewed at www.boreas.dk] BOREAS Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard 587 Fig. 2. Overview aerial orthophotographmosaic and fjord bathymetric data used tomap inner St. Jonsfjorden. The 2009 aerial images have a 0.5-m resolution while themarine bathymetry datawere griddedwith isometric grid cell size of 2 m.Locations of referenced dates fromprevious studies have been recalibrated and indicated. Subbottomprofiles are indicated bywhite lines 7 and 8 in the inner fjord, and are presented in Fig. 7. Coordinates are shown in metres, UTM zone 33N, WGS84. [Colour figure can be viewed at www.boreas.dk] 588 W esley R .F arnsw orth et al. B O R E A S logical and sedimentological mapping, as well as lithos- tratigraphical logging in the Piriepynten area (Figs 1, 2). Sea-floor data were collected in June 2013 (Fig. 2). Bathymetric data were collected on board UNIS R/V ‘Viking Explorer’ with a Kongsberg EM2040 multi- beam echosounder. Sea-floor sediment surface and subbottom acoustic sediment structures were studied with the Edgetech 2000 CSS combined side-scan sonar and subbottom profiler systems. The bathymetric data cover the entire St. Jonsfjorden up to the Osbornebreen and Konowbreen ice fronts. Nine samples of subfossil marine shells were selected fromthe southern flankof the inner fjordnearPiriepynten for radiocarbon dating, to constrain ages of glacial and marine events (Table 1). The samples were cleaned, weighed, photographed and subsequently sent for dating at the Radiocarbon Dating Laboratory at Lund Univer- sity and to theAngstr€om Laboratory at Uppsala Univer- sity. Ages are presented in calibrated kiloyears before present (OXCAL v4.2; Reimer et al. 2013). Samples were corrected for the marine reservoir age with 440 years (according to Mangerud & Gulliksen 1975; Mangerud et al. 2006). Metadata are presented in accordance with Hughes et al. (2016). The Quaternary geological maps were created in ArcMap 10.3 and the QPS Fledermaus Software package based on aerial orthoimagery from 2009 (NorwegianPolar Institute) andhigh-resolution sea- floor data from2013 (UNIS; Fig. 2). Previous studies and radiocarbon dates from the area presented by Forman (1989), Evans & Rea (2005) and Bunin (2015) are referenced to best constrain the glacial history of the fjord (Fig. 2). Results Terrestrial deposits, landforms and features We present a Quaternary geological map where the terrestrial landscape and sea floor of inner St. Jonsfjor- denhavebeenportrayed (Fig. 3). Themaphighlights the dominant components of the terrestrial landscape in the region: glacial, fluvial and periglacial deposits and landforms (Fig. 3). Glaciers cover approximately one- third of the mapped region and their catchments are controlledby topography.Ridgelinesof in situweathered bedrock and alpine terrain with shallow slope deposits divide the glacier systems. This terrain characterizes most of the higher elevations. The lower elevations are predominantly glacial deposits and glacifluvial fan sediments (where glaciers are land-terminating), and the flat-lying inter-catchment areas are dominated by solifluction material. Glacial sediments. – The most prominent deposits pre- sent in inner St. Jonsfjorden are Neoglacial glacial sediments. The glacial deposits are often confined to cirques, valleys and embayments, separated by ridgesTa bl e 1. L is to fr ad io ca rb on da te s fr om th is st ud y an d ot he rs ci te d he re in (a ll da te s co rr ec te d fo r re se rv oi r ag e 4 40 ye ar s) . L ab no . A ge (1 4 C a B P ) M at er ia l Se di m en ts L oc at io n E le va ti on (m a. s. l.) L at ./L on g. d1 3 C co rr . (& ) A ge (c al . a B P ; m ed ia n) A ge (c al . a B P ;1 r ) So ur ce (s am pl e da te ) L uS 10 79 0 18 00 5 0 M . ca lc ar ea T hr us te d si lt s P ir ie py nt en 10 78 0 3 0″ 55 N /1 30 04 ″1 3E 12 86 4 6 13 55 –1 18 1 T hi s st ud y (2 01 3) L uS 10 79 1 12 85 4 5 A . m on ta gu i T hr us te d si lt s P ir ie py nt en 8. 5 78 0 3 0″ 57 N /1 30 04 ″1 7E 75 8 57 83 1– 68 0 T hi s st ud y (2 01 3) L uS 10 79 2 52 30 5 0 L it ho th am ni on T hr us te d si lt s P ir ie py nt en 8 78 0 3 0″ 55 N /1 30 04 ″1 8E 55 17 6 8 56 08 –5 44 9 T hi s st ud y (2 01 3) L uS 10 79 3 16 15 6 0 M . ca lc ar ea T hr us te d si lt s P ir ie py nt en 10 78 0 3 0″ 56 N /1 30 04 ″1 3E 11 03 7 6 11 90 –9 64 T hi s st ud y (2 01 3) L uS 10 79 4 46 90 5 0 C . is la nd ic a T hr us te d si lt s P ir ie py nt en 10 78 0 3 0″ 54 N /1 30 04 ″5 4E 48 20 8 2 48 92 –4 78 4 T hi s st ud y (2 01 3) L uS 10 79 5 93 20 6 0 M . m od io lu s Sa nd + gr av el P ir ie py nt en 6 78 0 3 0″ 41 N /1 30 03 ″3 0E 10 00 4 12 2 10 10 9– 97 61 T hi s st ud y (2 01 3) L uS 10 79 6 90 70 5 5 M . tr un ca ta Sa nd + gr av el P ir ie py nt en 6 78 0 3 0″ 41 N /1 30 03 ″3 0E 95 96 6 0 97 36 –9 51 8 T hi s st ud y (2 01 3) U a- 52 52 4 92 41 3 1 M . m od io lu s Sa nd + gr av el P ir ie py nt en 5. 5 78 0 3 0″ 48 N /1 30 03 ″0 7E 0. 1 98 24 9 2 99 28 –9 68 8 T hi s st ud y (2 01 5) U a- 52 52 5 91 89 3 8 M . tr un ca ta Sa nd + gr av el P ir ie py nt en 5. 25 78 0 3 0″ 48 N /1 30 03 ″0 7E 1. 2 97 30 8 9 99 00 –9 59 5 T hi s st ud y (2 01 5) D IC 30 56 86 90 8 5 M . tr un ca ta Su b- lit .S an d L øv lie br ee n 3. 5 0 92 31 1 17 94 37 –9 02 1 F or m an (1 98 9) D IC 30 55 99 60 9 0 M . tr un ca ta Su b- lit .S an d A nk er br ee n 13 .5 0 10 85 9 16 0 11 14 9– 10 58 1 F or m an (1 98 9) G U 80 69 92 0 60 M . tr un ca ta T hr us te d si lt s P ir ie py nt en 21 1. 6 51 8 65 56 7– 42 8 E va ns & R ea (2 00 5) G U 80 70 98 40 7 0 M . tr un ca ta G ra ve lb en ch G je rt se no dd en 19 .5 2. 2 10 63 2 12 3 10 79 4– 10 41 3 E va ns & R ea (2 00 5) G U 80 72 10 13 0 90 M . tr un ca ta F or es et sa nd s B ul lb re en 15 .5 1. 6 11 05 2 14 7 11 24 3– 10 76 0 E va ns & R ea (2 00 5) G U 80 73 10 48 0 90 M . tr un ca ta M ar in e si lt s G je rt se no dd en 13 2 11 57 0 18 5 11 84 1– 11 25 6 E va ns & R ea (2 00 5) G U 80 74 10 35 0 80 M . tr un ca ta D ia m ic t B ul lb re en 12 2. 2 11 36 0 13 9 11 64 3– 11 19 7 E va ns & R ea (2 00 5) G U 80 75 10 27 0 80 M . tr un ca ta D ia m ic ts ur fa ce C op pe r ca m p (E ) 19 2 11 25 1 12 1 11 50 2– 11 09 4 E va ns & R ea (2 00 5) H H 12 .9 56 11 34 5 10 2 N . la br ad or ic a M ar in e m ud s O ut er F jo rd 1 11 78 0 3 1″ 79 N /1 20 27 ″0 8E 12 80 5 95 13 02 9– 12 67 8 B un in (2 01 5) BOREAS Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard 589 Fig. 3. OverviewQuaternarygeologicalmapof inner St. Jonsfjorden.Cross-cutting glacial deposits record a complex anddynamichistory.Coordinates are shown inmetres,UTMzone 33N,WGS84. [Colour figure can be viewed at www.boreas.dk] 590 W esley R .F arnsw orth et al. B O R E A S Fig. 4. DetailedQuaternarygeological map of Piriepynten, inner St. Jonsfjorden. Thermophilousmolluscs identified in earlyHolocene beach sediments just to the south of Piriepynten (white star). Coordinates are shown in metres, UTM zone 33N, WGS84. [Colour figure can be viewed at www.boreas.dk] B O R E A S D ynam ic H olocene glacialhistory ofS t.Jonsfjorden,S valbard 591 and alpine terrain. An extensive but subtler glacial deposit is visible along the flanks of the inner fjord starting at Piriepynten and tracing back to the modern- day ice fronts of Konowbreen and Osbornbreen. Here, we introduce and describe this deposit, but the compo- sition and age will be presented in the section Coastal moraine. The glacial deposit is situated along the coast at the northernmost point of Piriepynten and a submarine ridge can be traced on the bathymetrical data set across the fjord linking to similar sediments on the northern side of the fjord. This is a deposit of low- crested glacial sediments that flanks the coast. The sediments outline a glacial system composed of the considerably expanded marine terminating glaciers Konowbreen and Osbornebreen (Fig. 3). The fjord-parallel till drape of the expanded Konow- breen and Osbornebreen has been cross-cut by the smaller tributary valley and cirque glaciers inside of Piriepynten. The subsequent advance of a tributary glacier is perceptible at the snout of Smalgangen on the northern side of St. Jonsfjorden where local glacial deposits cross the low-elevation, fjord-parallel till drape (Fig. 3). The subsequent advance of a tributaryglacier is also visible where the ice-marginal deposits of Charles- breen cut through coast-parallel glacial deposits (Fig. 3). The western contact can also be seen in detail as the cross-cutting relationship between an extended Osbornebreen and a subsequent extension of Charles- breen (Fig. 4). A glacier advance subsequent to themax- imum extension of Charlesbreen is indicated, where the cirque glacier Anna Sofiabreen seems to have advanced into and deformed Charlesbreen lateral moraine. The glacial sediments and their characteristics can be better visualized in the detailed Quaternary geological map from the Piriepynten area where glacial deposits originate from five different glaciers (Løvliebreen, Gunnarbreen, Anna Sofiebreen, Charlesbreen and Osbornebreen; Fig. 4). Some glacial deposits are sparsely vegetated and appear relatively stable, while other deposits exhibit widespread slumping, fluvial-erosional cutting and dead-ice melt-out, making them unfavour- able for vegetation growth (Fig. 4). Lake sediments. – Lake sediments have been identified and mapped on the eastern glacifluvial fan system flowing from Gunnarbreen (Fig. 4). These silty lake deposits are partially preserved between fluvial channels on the western flankof the fan system where the braided river system is less active today. The lacustrine sediments are located on the south side of the glacial deposits flanking the coast. Deposits are estimated to reach decimetre thickness and are not widespread. Raised marine sediments. – Raised marine deposits are highlighted in the detailed map of Piriepynten (Fig. 4). These sediments are not common in the inner St. Jonsfjorden area and the highest marine deposits found in the area are located just to the west of the Løvliebreen marginal moraine at ~20 m a.s.l. (Fig. 4). The upper Holocenemarine limit is believed to be ~45 m a.s.l. in the outer fjord andas lowas 30 mnearLøvilebreen (Forman 1989; Evans&Rea 2005).Nomarine deposits or features above 20 m elevation were identified in the Piriepynten area. Themost prominent deposit ofmarine sediments is preserved south of the bedrock knob at Piriepynten. The raised beach deposits have minor vegetation cover and consist of sandy gravels and cobbles predominantly disc or plate formed with subrounded to subangular mor- phology. Several sets of raised beach ridges are preserved on the surface of the marine sediments. The uppermost beach ridges are vegetated and display polygonal struc- tures with sorted and unsorted patterned ground caused by periglacial sorting at the surface. The upper terrace of Fig. 5. Lithostratigraphical log of early Holocene sediments exposed near Piriepynten, St. Jonsfjorden. Coarsening upwards sequence overlying silty matrix-supported diamict. Calibrated radiocarbon dated samples indicated with ages given in years BP. [Colour figure can be viewed at www.boreas.dk] 592 Wesley R. Farnsworth et al. BOREAS this deposit is ~8 m a.s.l. and forms the natural exposure on thesouthernhalfof thesediments,whichare estimated to be at least 5 m thick. The upper 2.5 m of the section is characterized as a coarsening upward sequence that has been divided into two main units (Fig. 5). • Unit 1: in the lowermostpart there ismore than30 cm of massive, silty-sandy, matrix-supported diamict with outsized pebbles and cobbles frequent in its lower part. Based on the sediment composition we interpret this as a morainal bank sediment (Powell & Domack 1995), deposited at shallowdepth in front of a retreating ice front. • Unit 2: bove Unit 1 is 2.2 m of stratified sorted sediments that coarsenupwards. In the lower part the unitconsistsofasiltysandsequencewithcrudebedsof sand and in situ Mya truncata shells (Fig. 5), which conforms to planar-bedded (cm scale) silty togravelly sands. Several half bivalve Modiolus modiolus shells and fragments of blue shell (Mytilus edulis) were identified (Fig. 5). These shells were not in living position but do not show signs of long transport. The unit coarsens upwards to cobbles and gravels. In some beds there is silt and sand,while in other coarser beds, fine-grainedsedimentsarescarce.Thetop0.25 misclast- supported coarse sand, gravel, and cobble with imbrica- tion structures. The clasts are plate formed with sub- rounded to subangular morphology. Based on the sediment composition we interpret this as a regressional sublittoral sequence, capped by beach sediments. Numerous shells were sampled and identified in the section, and two shells have been dated (Fig. 5). The lowermost M. modiolus shell and the paired in situ M. truncata are dated to 9.820.09 and 9.730.09 cal. ka BP, respectively (Fig. 5; Table 1). Although the lower shell (M. truncata) gives a slightly older age, the dates overlap and the M. truncata could have burrowed beneath the M. modiolus. These ages correspond well with aModiolus modiolus, which was sampled at the foot of thenatural section in2013,dating to10.00.12 cal. ka BP (Table 1). This provides the earliest shell age effec- tively constraining the minimum age of deglaciation of inner St. Jonsfjorden (Table 1; Sample LuS-10795) as well as constraining the timing of relative sea level at ~8 m a.s.l. Coastal moraine. – To the east of Piriepynten, a low- crested moraine was identified (Fig. 4). The landform has little to no vegetation or lichens and generally lacks cobbles and larger clasts on its surface. The surface is dominated by dark clayey-silt with abundant mollusc shells. A light colouredprecipitatewas visible on some of thedesiccatedmud-cracked surfaces. In anatural coastal section, the silty sediments overlie crudely stratified sorted sands and gravels with occasional cobbles inter- mixed. A compact matrix-supported silty diamict with stratified sands and gravels overlies the sorted sands and gravels. The sediments fine upwards into a shell-rich silty clay that is visible on the surface. We interpret this as Fig. 6. A. Beach sediments at Piriepynten with view towards Osbornebreen to the NE. The over-deepened drainage channel cuts the early Holocene beach sedimentswith the spillway to the bottom-left of the image.Moraine systemsmarked in background. B, C.Mollusc shells picked from the stratigraphical section in the beach sediments, paired in situMya truncata andModiolusmodiolus, respectively.D. Bedrockexposurewith cross-cutting striae. Finer, widespread striaewith near north–south orientation, characterized as local flow. Larger, less prevalent striaewith east– northeast to west–southwest orientation, regarded as regional, fjord-parallel flow. [Colour figure can be viewed at www.boreas.dk] BOREAS Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard 593 594 Wesley R. Farnsworth et al. BOREAS marine sediments that have been pushed up onto the beach by an advancing glacier. Shells of numerous species were sampled and identi- fied fromthe silty sediments includingMacomacalcarea, Astarte montagui, Lithothamnion sp., Chlamys islandica and Mya truncata. Six samples were dated to con- strain the age of the moraine. The youngest age from the moraine is 0.760.06 cal. ka BP, suggesting that the feature was deposited sometime after 1200 CE (Table 1; Sample LuS-10791). The five other shell dates from the coastalmoraine range from9.60.06 to1.10.08 cal. ka BP suggesting conditions in the inner fjord must have beensufficiently freeof ice toallowthesemolluscs toexist (Table 1). Crevasse-squeeze ridges. – Fragmented networks of cre- vasse-squeeze ridges (CSRs) have been identified and mapped in the foreland of Gunnarbreen and Charles- breen(Fig. 4).TheCSRsareonlypresentproximal to the wide marginal moraines (Fig. 4). The landforms have low preservation potential and are often reworked by meltwater and dead-ice melt (Lovell et al. 2015b). The ridges consist of matrix-supported diamict sediments andareoftenpreserved ingeometricnetworks locatedon raised till plains protected from meltwater erosion (Schomackeret al.2014).The landformsarealsopresent in front of Løvliebreen and numerous other glaciers located in St. Jonsfjorden although not shown in Fig. 3 (Farnsworth et al. 2016). Drainage channels. – Several small drainage channels cut through the southern part of the raised marine sediments.Thechannels run to the southwest and initiate at the crest of the beach ridges around 8 m a.s.l. (Figs 4, 6A).Thechannels cutdownnearly2 mandat the floorof the channels, outsized, striated clasts along with shells (Mya truncata) have accumulated. We refer to these features as the initiation drainage channels (Fig. 6B). The small channels are cut by a large relict drainage channel that runs approximately east to west. This channel is presently comprised of two small elongated ponds that have been divided by subsequent glacifluvial sedimentation from Gunnarbreen (Figs 4, 6A). This east–west trending channel has cut through the marine sediments and has over-deepened and formed the present-day basins. The channel is roughly 250 m long from east to west, up to 7 m deep and nearly 10 m wide. The surface level of thewater in the basinswas calculated to be ~3 m a.s.l. and the water is up to ~2 m deep in the large pool (Fig. 6A). Striated boulders of 50–100 cm in size are also present in the over-deepened channel. The natural section on the northern flank of this channel is Fig. 8. Timeline of the Holocene glacial histories from fjords on the west coast of Spitsbergen. Ages are presented in calibrated kiloyears (ka) BP andcorrected formarine reservoir age. St. Jonsfjorden 14Cages and results are presented above the timeline,while Isfjorden,VanMijenfjorden and VanKeulenfjorden are presented below.Dark arrowheads constrain events, while hollowed arrowheads present ages that oncewere constraining. ShadedModiolusmodiolusandwhitebivalve symbols represent dated samples from theSt. Jonsfjorden section (Fig. 5B,C).Blackbivalve symbols indicate dated shells, sampled from the coastal moraine ridge.[Colour figure can be viewed at www.boreas.dk] Fig. 7. Bathymetry data and subbottom profiles (SBFs) of inner St. Jonsfjorden. A. Bathymetry data of inner St. Jonsfjorden with SBFs 7 and 8 indicatedbywhite lines.B.DetailedQuaternarygeologicalmapof theseafloor in innerSt. Jonsfjorden.C,D.Zoominonthebathymetryof the inner fjord, with clear examples of cross-cutting glacial landforms and deposits. E. Amodified SBFacross the 2nd order moraine (blue colour, Fig. 7B), oblique tothestreamlinedfeaturesandrecessionalmoraines (upward,whitearrows).F.ASBFdistal to the2ndordermoraine, showingstreamlined landforms and crevasse squeeze ridges (bothmarked by upward, black arrows)with infill and drape of/by acoustically stratified sediment (marked by downward, black arrows). Coordinates are shown in metres, UTM zone 33N, WGS84. [Colour figure can be viewed at www.boreas.dk] BOREAS Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard 595 where theM. modiolusmolluscswere sampled (Fig. 6C). We refer to this as theover-deepeneddrainage channel or main spillway. A final northern, significantly smaller drainage chan- nel is also present. This drainage channel flows to the northwest and runs between the Piriepynten bedrock knob and the coast-parallel moraines just north of the raised marine terraces. Although similar in length, this drainage incision is~2–3 mdeep (Fig. 4).Werefer to this as the second spillway. Glacial striations . – Abedrock exposure to the south of the over-deepened channel displays cross-cutting glacial striations (Fig. 6D). One set of striae is orientated approximately east-northeast to west-southwest (~60– 240°).The striations are coarse andweathered.Theother set of striae runs nearly north to south (~0–180°). These striae are more widespread, rough and thinner than the first set. They are present on bedrock convexities and prominences, while the fjord-parallel striations are present in the concavities (Fig. 6D). Fjord bathymetry and submarine geomorphology The bathymetric data of St. Jonsfjorden reveal bedrock structures and sediment basins as well as glacial land- forms (Figs 2, 3). In the overviewmap of the inner fjord, the most prominent feature is a 3-km-wide north–south trending shallow sill initiating from Piriepynten. We focus mainly on the glacial landforms preserved in the inner fjord to the northeast of this sill. Parts of the inner basin reach nearly 150 m below sea level. Moraines. – From the bathymetric data set, numerous large lobate ridges have been identified on the sea floor. The landforms are arcuate ridges spanning 1 to 2 km in length and 0.5 to 1.5 km in width with prominence ~25 m. This type of ridge occurs distal to the present front of Gaffelbreen, Konowbreen and Charlesbreen and the shape and orientation reflect the present-day glacier fronts (mapped as 2nd order moraines; Fig. 3). A significantly larger lobate ridge occurs in the middle of fjord, trending north–south Fig. 9. Oblique aerial photograph modified from the Norwegian Polar Institute looking south onto Piriepynten taken in 1936. Site locations identified aswell as roughoutlines ofLIAmaximumpositions forGunnarbreen,AnnaSofiabreenandCharlesbreenmarked bywhite dotted lines. Green dotted line marks the extent of the Piriepynten moraine (coastal moraine) sometime in the last 500 years. Arrows represent cross-cutting relationshipsbetweenCharlesbreenandKonowbreen-Osbornbreen sometimeduring the last 500 years aswell asGunnarbreen and themain fjord outlet glacier during the earlyHolocene.Note thatGunnarbreenwas at least 750 mmore extended during the earlyHolocene and terminated into the fjord system. [Colour figure can be viewed at www.boreas.dk] 596 Wesley R. Farnsworth et al. BOREAS corresponding with the location of the prominent sill north of Piriepynten (1st order moraine). The ridge is over 2 km long and roughly 1.5 km wide. The crest of the ridge is prominent and rises up from basins >70 m to roughly 30 m below sea surface. The smaller ridges (2nd order moraines), based on their proximity to the glaciers, shape and relief, are interpreted as terminal moraines that originate from Gaffelbreen, Konow- breen and Charlesbreen. The largest lobate ridge is interpreted as a terminal moraine deposited from a larger confluent Konowbreen-Osbornebreen. We refer to this ridge as the Piriepynten (1st order) moraine. Based on the spatial occurrence of these landforms, we suggest the Charlesbreen and Konowbreen (2nd order) moraines formed subsequent to the larger Piriepynten moraine. We are unable to constrain the age of the Gaffelbreen moraine relative to the constrained Piri- epynten moraine. Debrisflow aprons. – Aprons of debrisflow deposits are observed on the sea floor distal to the large moraines. The largest apron is located on the distal side of the large Piriepynten moraine extending ~3 km from the moraine crest and filling the width of the fjord basin (>2 km).The surfaces of the aprons are non-uniform, which is probably caused by the irregular distribution of iceberg scours, slumps and slides on the surface. The shape and irregular surface leads to an interpretation of the deposit as an accumulation of debrisflow lobes forming an apron of sediments. Similar submarine massflow deposits have been described from surge-type glaciers in Svalbard (Ottesen et al. 2008; Flink et al. 2015). Streamlined features. – Widespread streamlined fea- tures have been mapped in inner St. Jonsfjorden and are interpreted as glacial lineations (Ely et al. 2016). They appear in subparallel swarms and are most promi- nent between the Piriepynten moraine and the present- day ice front of the marine terminating glaciers Konow- breen and Osbornebreen (Fig. 3). The streamlined feature orientations vary from north–south to nearly east–west, depending on the alignment of the fjord basin (Figs 3, 7A,B).Mapped lineations in someplaces extend over a kilometre in length and mostly exceed 5 m relief, ranging 1–15 m in height (Fig. 7). The features are roughly 10 m wide and spacing between crests ranges from 50 to 500 m (averaging ~200 m). Some of the streamlined features are overlain by minor moraines located at thewesternbranchof the inner fjord (Fig. 7B– D). Transverse ridges. – Ridge segments transverse to gla- cier flow are present throughout inner St. Jonsfjorden (Fig. 7A, B, C, E). The ridges are up to 1.5 km in length and of relatively low relief (~5 m). The ridges are located sporadically throughout the inner fjordandare common and frequently spaced in the inner-most embayments corresponding to Osbornebreen and Konowbreen. Inter-ridge spacing ranges from 50 to 400 mwith typical distances from crest to crest of ~150 m (Fig. 7A, B). The ridges are less linear than the glacial lineations and other streamlined features and canbe foundoverprinting these features. Ridges of similar appearance, morphology and spacing havebeen identified in other Svalbard fjords and have been interpreted as retreat moraines (Flink et al. 2015, 2017). Crevasse-squeeze ridges. – A geometric ridge network has been mapped and interpreted as crevasse-squeeze ridges in inner St. Jonsfjorden (Lovell et al. 2015b; Figs 3, 7). The rhombohedral structures are present on a flat basin along with streamlined glacial features (Fig. 7D). This ridge network is predominantly oblique to the streamlined features. The crevasse-squeeze ridges are low-relief features (<5 m). The eastern part of subbot- tom profile eight crosses over the network of ridges (Fig. 7D, F). The profile shows numerous parabolas interpreted to be ridges orientated obliquely or perpen- dicular to the profile. In some inter-ridge areas several metres of sediment seem to have accumulated based on the flat-lying, horizontal strata between topographical features (parabolas; Fig. 7). This suggests that the landforms have been draped by and in some cases masked completely by marine sediments subsequent to their formation. Discussion St. Jonsfjorden deglaciation, Atlantic water and late Holocene glacial history The outer parts of the major fjords on the central west coast of Spitsbergen were deglaciated at a similar period in time between 12.5 and 11.7 cal. ka BP (Fig. 8; Mangerud et al. 1992; Forwick & Vorren 2009; Kempf et al. 2013; Bunin 2015). Although the exact age of deglaciation is unknown, the minimum constraining deglaciation age for the outer St. Jonsfjorden is 12.80.10 cal. ka BP (Bunin 2015). This compares well with the range of the earliest dates of foraminifera and shells collected from the mouths of Isfjorden, Northern Bellsund, Van Mijenfjorden and Van Keulenfjorden dating to ~12.5, 12.7, 11.7 and 11.8 cal. ka BP, respec- tively (Landvik & Salvigsen 1987; Mangerud et al. 1992; Bratlie 1994; Forwick & Vorren 2009). Prominent bedrock thresholds at the western extent of Van Mijen- fjordenandVanKeulenfjordenmayhave contributed toa later deglaciation of these fjords (Kempf et al. 2013). Kongsfjorden north of St. Jonsfjorden, a large trough system without a prominent bedrock threshold at the mouth, is believed to have deglaciated between 16.6 and 14.4 cal. ka BP (Lehman & Forman 1992; Henriksen et al. 2014). BOREAS Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard 597 Cross-cutting striations on bedrock suggest a north– south orientated ice flow subsequent to fjord-parallel ice flow in St. Jonsfjorden. Stratigraphical evidence proxi- mal to the striated bedrock suggests the north–south ice- flowphasemustpre-date9.8 cal. kaBP(Fig. 6).At some point between the deglaciation of outer St. Jonsfjorden (12.80.10 cal. ka BP) and before the first molluscs migrate into the inner fjord system (~10.0 cal. ka BP), Gunnarbreen advanced to a more extensive position. Although the full glacier extent is unknown, the snout extended at least 750 m further than it did in 1936 (Figs 3, 6D, 8, 9). The HTO is fingerprinted by the migration and presence of thermophilous molluscs, and warm water foraminifera to waters around Svalbard between 11.2 and 7.7 cal. ka BP (Salvigsen 2002; Hald et al. 2004). Many of these species only lived in the waters around Svalbard during this short-termwarm period in the early Holocene (Fig. 8). At present, many of these species can be found along the coast of mainlandNorway and to the east along the coast of Russia (Salvigsen 2002; Deni- senko et al. 2007). Few M. modiolus shells have been identified on Svalbard and even fewer have been radiocarbon dated. In a review of thermophilous mol- luscs from Svalbard, Salvigsen et al. (1992) highlights several dated individuals from Isfjorden (Fig. 8). There is also a reference to identification of a M. modiolus in St. Jonsfjorden, but dating was performed on a Mya truncata (DIC 3056; Table 1) in the same stratigraphical position (Forman 1989; Salvigsen et al. 1992). These dates suggest the species inhabited the waters around Svalbard from ~10.0 to 7.4 cal. ka BP. The two new ages ofM. modiolus from Piriepynten presented in this study corroborate well with the earlier range of ages accumu- lated from previous Svalbard studies (Fig. 8; Salvigsen et al. 1992; Hald et al. 2004; Beierlein et al. 2015). Not only is thismollusc age interestingwith regards to the early migration of the M. modiolus species to Svalbard, the sample also provides the constraining age for the deglaciation of the inner part of St. Jonsfjorden (Fig. 8). This suggests that relatively warm North Atlantic waters were present at or immediately after the deglaciation of the inner fjord. Studies on mollusc shell growth bands and shell microstructures suggest strong seasonalitywith short favourable summers and long cold winters on Svalbard (Salvigsen et al. 1992; Beierlein et al. 2015). Studies of modern processes and glacier response to increased ocean temperatures on Svalbard (Bartholomaus et al. 2013) indicate high sensitivity where a minor change in subsurface water temperatures Fig. 10. Generalized map highlighting a simplified sequence of glacier front positions (1–6) in inner St. Jonsfjorden since ~500 years BP (0.520.07 cal. ka BP; GU 8069). Stage 7 is constrained by submarine morphology and a historical oblique aerial image from 1936 indicates a heavily crevassed Konowbreen in the early phase of a surge event (TopoSvalbard 2017). [Colour figure can be viewed at www.boreas.dk] 598 Wesley R. Farnsworth et al. BOREAS can have a large impact on glacier conditions (Luckman et al. 2015). The contemporary understanding of the modernprocesses and the constraining deglaciation ages provided by the M. modiolus shells in St. Jonsfjorden favour warm Atlantic waters as a driving factor for the deglaciation of the area. Themid to lateHolocene on Svalbard is characterized by decreasing temperatures in regional waters, which is also seen in St. Jonsfjorden (Hald et al. 2004; Bunin 2015). Bunin (2015) interprets cooling conditions from distal glaciomarine sediments dating ~4.0 cal. ka BP, and highlights an increase in glacial activity interpreted from the fjord sediments at ~1.7 cal. kaBP.Based on our shell dates from the coastal moraine, we suggest the inner fjordwas ice-free enough to allowmolluscs to exist between 10.0 and 0.76 cal. ka BP (Table 1; Fig. 8). The youngest shell date we present (0.760.06 cal. ka BP) corroborates with previous work that constrained the Piriepynten moraine to a late Holocene glacial advance (0.520.07 cal. ka BP; Evans&Rea 2005). Based on the ages of subfossil shells we suggest the Osbornbreen coastal moraine was deposited sometime in the last 500 years (Figs 8, 9; Evans & Rea 2005). In the late Holocene, numerous advances have been recorded or dated from other large glacier systems in both Van Mijenfjorden and Van Keulenfjorden (Hald et al. 2004; Kristensen et al. 2009). These events have been sug- gested to be related to surge-type behaviour (Ottesen et al. 2008; Kristensen & Benn 2012). Based on the network of crevasse-squeeze ridges identified at the sea floor in front of the Konowbreen-Osbornebreen glacier system (corresponding to the Piriepynten moraine), we suggest this advance may also be related to surge-type behaviour (Fig. 7A, B). However, we note that climatic conditions around the time of the glacier advance were also favourable for a positive glacier mass balance (Fig. 8). A conceptual model for Neoglacial glacier dynamics in St. Jonsfjorden By mapping the outermost extents of the larger glaciers in the inner fjord, numerous cross-cutting relationships appear (Figs 3, 9, 10). The outermost extent ofOsborne- breen at Piriepynten is constrained to 0.520.07 cal. ka BP (GU 8069; Stage 1 in Fig. 10; Evans & Rea 2005). Subsequentadvancesarenumbered2–7 (Fig. 10). Stages 4 and 7 do not correspond to advances, but rather ice- front positions interpreted from the Isachsen Spitsber- gen Expedition (Hoel et al. 1915–1917) and from bathymetry data combined with the oblique 1936 aerial imagery provided by theNorwegian Polar Institute. The innermostpositionofOsbornebreen indicates theglacier front position in 2009. Aconceptualmodelhasbeendevelopedonthebasisof the series of glacial advances in the inner St. Jonsfjorden (Fig. 11). The concept is that the local glaciers exhibit small advances due to unbalanced glaciodynamic con- ditions, and not in direct response to climate. If a tributaryglacier is at equilibrium and has a stable profile given theback-stress causedbytheprimary trunkglacier, retreat of the trunk glacier will perturb the balance of the tributary. This leads to a re-organization of the over- steepened glacier profile, which is glaciodynamically expressed as an advance. The tributary then advances until it reaches a new dynamic equilibrium state that is marked by a new quasi-stable marginal position due to the lack of back-stress (Fig. 11; Wuite et al. 2015). The re-organization of a glacier profile related to a decrease in back-stress can be a combination of the Fig. 11. Conceptual model introducing a dynamic advance resulting from unbalanced equilibrium leading to profile re-organization. A. Trunkglacier is inbalance, and tributaryglacierprofile is in equilibrium with the back-stress of the trunk. B. The retreat of the trunk glacier creates an unbalanced system. C. Tributary glacier re-organizes its surface profile in response to the release of back-stress which results in an unsustainable advance of the ice-front (Scambos et al. 2004; Wuite et al.2015).D.Tributaryglacier retreatsback toanewequilibriumstate given the new conditions. Tributary pursues a newbalanced profile.We choose to exclude a sea-level component from our conceptual model, yet identify the importance of further addressing this aspect. [Colour figure can be viewed at www.boreas.dk] BOREAS Dynamic Holocene glacial history of St. Jonsfjorden, Svalbard 599 switching of thermal regime, or merely the shedding of accumulatedmass that has been gained over an extended period of time. Furthermore, this behaviour is applicable at awide range of glacier scales – from the small cirque of Anna Sofiebreen that advances into the Charlesbreen basin (Figs 4, 9), up to the size of the collapse of the Larsen B ice shelf in Antarctica and the subsequent advancesof thenumerous tributaries (Scamboset al.2004; Wuite et al. 2015). Although the degree towhich glaciody- namics vary between marine and terrestrial terminating glacierswith regards to back-stress is not well understood, we assume the response to be amplified in the marine environment (Nicket al.2010).Wechoose toexcludeasea- level component from our conceptual model, yet identify the importance of further addressing this aspect. This conceptual model provides a potential answer to the numerous descriptions of glacier advances on Svalbard that correspond to the periods of unfavourable climate conditions for glacier growth in both terrestrial and marine environments (Lønne 2005; van der Bilt et al. 2015). The model could as well be extended to the early Holocene advance of Gunnarbreen into a shallow marine environment subsequent to the retreat of the dominant tidewater glacier in the fjord (Fig. 9). The model also highlights how different glaciers may respond uniquely to an external force. Thus, the model steps away from the paradigm that glacier advances fit either to a phase of increased winter precipitation or reduced summer temperatures (Solomina et al. 2015). It is reasonable to expect increasedprecipitationdue to warmersea-surface temperaturesandreducedfjordcover (by sea ice or glacier), but this is not likely to drive glacier advances of that order of magnitude (Bintanja & Selten 2014). Thus, warm fjord-waters resulting in an increase in precipitation may well have occurred, but are not necessarily the sole driver of the glacier advances. Fewer assumptions are made interpreting these small advances to be dynamically driven. Additionally, the time-trans- gressive nature of the events suggests that the local advances are responses to glacier-specific conditions (i.e. profile re-organization) and not a regional amplification of precipitation. The principle of parsimony suggests fewer assumptions are taken for a dynamic advance as opposed tonumerousclimate-drivenadvances foreachof the tributaryglaciers, bothmarineand terrestrial, located in inner St. Jonsfjorden. The Svalbard region today is a hotspot for dynamic glacier behaviour that may not be directly in phase with atmospheric or oceanic climatic forcings (Jiskoot et al. 2000; Sevestre & Benn 2015). We suggest that this may be a key to the past, and the same processes were occurring at other periods during the Holocene. Piriepynten landscape development through the Holocene The onshore landscape and the sea-floor bathymetry of St. Jonsfjorden are dominantly fingerprinted by glaciers during the later half of the Holocene, driven by the Neoglacial –LIAglacier expansion.Thewidespreadand prominent Neoglacial glacial deposits and the subse- quent fluvial runoff greatly mask, deform and rework previous sediments, landforms and features (Landvik et al. 2014). In some locations early Holocene deposits and features are visible, although subtle (Lønne 2005). The older deposits are mostly visible where younger events have cut, cleared or reworked them, exposing a ‘window’ into the past environment. The main drainage channel that scoured fromeast towardswest just southof the Piriepynten bedrock knob has opened the window into the early Holocene (Fig. 4). We suggest that the glacier advance of the confluent Osbornebreen-Konowbreen that formed the Piriepyn- ten moraine (sometime in the last 500 years) made identification of the early Holocene strata possible. Based on our interpretation of the landscape, we suggest an ice-dammed lake formed as a result of the meltwater fromGunnarbreen not being able to enter the fjord. The coastal moraine and the extended Osbornebreen, which pushed the sediments from the basin on the coast, trapped the meltwater from Gunnarbreen, which sub- sequently rose up to the raised marine terraces to the southeast of Piriepynten. After several small initiation drainages, the main drainage burst and drained the lake to the west around Piriepynten. Although this event has not been directly documented, an ice-marginal lake serves as a plausible explanation for the formation of the drainages through the raised marine sediments. This channel washed the sediments from the striated bedrock as well as cut through the raised marine sediments where the two M. modiolus molluscs were identified (Fig. 6C, D). Without this exposed stratigra- phyandglacial striae, the earlyHolocene glacier advance from Gunnarbreen would be impossible to detect, despite the fact that the glacier extended further than during the LIA maximum (Fig. 9). This confirms the concept of Landvik et al. (2014), which warns that our reconstructions arebiased towards the younger events. It also highlights the challenge of identifying the often subtle signatures of early Holocene glacier events. Conclusions We present combined terrestrial andmarine data from a small fjord on the west coast of Spitsbergen that exhibits an array of varying climatic conditions and glacier fluctuations throughout what was previously believed to be anuneventful Holocene. These data provide a holistic view of the complexity and the dynamics of these environments. • An early Holocene advance of the St. Jonsfjor- den glacier, Gunnarbreen, is constrained between 12.80.10 and 10.00.12 cal. ka BP corresponding to the Holocene maximum glacier extent. 600 Wesley R. Farnsworth et al. BOREAS • Identification and 14C dating of the thermophilous bivalve molluscModiolus modiolus to 10.00.12 cal. ka BP suggests a rapid northward migration of the species shortly after deglaciation, suggesting an early Holocene invasion of warm water as a potential driving factor of ice retreat. • We present insight into the onset, climax and break- downof theNeoglacial-Little IceAgemaximuminSt. Jonsfjorden. The Osbornebreen terminus is currently located over 8.5 km up-fjord from its Neoglacial maximum extent and cross-cutting relations suggest numerous subsequent advances of smaller tributary glaciers following the retreat of Osbornebreen. • We suggest that late Holocene glaciodynamic beha- viour of tributary glaciers in St. Jonsfjorden, follow- ing the retreat of the trunk glacier, may be related to a dynamic re-organization of their profiles. Acknowledgements. – The authors acknowledge financial support from the University Centre in Svalbard Research Fund, which aided in costs of fieldwork and radiocarbon dating. Lena Rubensdotter kindly advised us inmap preparation. Graduate students participating in theUNIS coursesAG-332andAG-832 in 2013–2015are thanked for field assistance. The authors acknowledge Astrid Lysa and Tom Bradwell for their constructive reviews as well as Jan A. Piotrowski for editorial support. References Alsos, I. G., Sj€ogren, P., Edwards, M. E., Landvik, J. 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Aerial image of Pettersenbreen foreland, southeastern Edgøya (Norwegian Polar Institute). Over 400 previously undocumented Svalbard surge-type glaciers identified Wesley R. Farnsworth a,b,⁎, Ólafur Ingólfsson a,c, Michael Retelle a,d, Anders Schomacker b a Department of Arctic Geology, University Centre in Svalbard (UNIS), N-9171 Longyearbyen, Norway b Department of Geology, UiT The Arctic University of Norway, N-9037 Tromsø, Norway c Faculty of Earth Sciences, University of Iceland, Sturlugata 7, IS-101 Reykjavík, Iceland d Department of Geology, Bates College, Lewiston, ME 04240, USA a b s t r a c ta r t i c l e i n f o Article history: Received 1 December 2015 Received in revised form 21 March 2016 Accepted 22 March 2016 Available online 13 April 2016 Identifying glaciers that exhibit surge-type behavior is important when using evidence of ice front fluctuations as a proxy for reconstructing past climate oscillations. This study identifies previously undocumented surge-type glaciers in Svalbard, based on the presence of crevasse squeeze ridges in glacier forelands. Crevasse squeeze ridges are landforms suggested to be unique to surging glacier land systems. Estimates vary greatly as to the actual percentage of surge-type glaciers in Svalbard, and consequently their distribution pattern is poorly under- stood. A detailed survey of recent (2008–2012), high-resolution aerial imagery from TopoSvalbard, provided by the Norwegian Polar Institute, allowed for a survey of all the glacier forelands in Svalbard. Before our study, 277 individual glaciers in Svalbard have been documented to exhibit surge behavior. By using crevasse squeeze ridges as indicators of surge behavior, we have identified 431 additional glaciers that have surged. We suggest that this is amodest value as the unique surge landformswere not visible in approximately one-third of the fore- lands with documented surge histories. Limits to the crevasse squeeze ridge technique are presented and poten- tial controlling factors for crevasse squeeze ridge formation/preservation are discussed. © 2016 Elsevier B.V. All rights reserved. Keywords: Crevasse squeeze ridges Svalbard Glacier surge Aerial imagery 1. Introduction Holocene climate has fluctuated between warmer and cooler periods as well as relatively humid and dry conditions on decadal to centennial timescales (Wanner et al., 2011). ‘Polar amplification’ refers to the sensi- tivity of Arctic regions to these climatic fluctuations and how the Arctic areas are affected by relatively small shifts in temperature and precipita- tion (Masson-Delmotte et al., 2013). Arctic glaciers and ice caps respond to even small changes in winter precipitation and summer temperatures (Masson-Delmotte et al., 2013), making them valuable indicators of changes in climate (Oerlemans, 2005). Consequently, the reconstruction of Arctic glacial oscillations and chronologies through the Holocene provides valuable insight into antecedent climate scenarios. Understand- ing that individual glaciers may respond uniquely to shifts in climate is vital (Røthe et al., 2015). Furthermore, surge behavior complicates the matter as glacier fluctuations sometimes are driven by internal ice dynamics (surges) that some have been suggested to be unrelated to climate (Meier and Post, 1969; Yde and Paasche, 2010). It is therefore important to establish if ice-front oscillations for individual glaciers reflect mass-balance shifts controlled by climate, variations in internal dynamics (surge activity), or a combination of the two. The aim of this study is to identify previously undocumented surge- type glaciers on Svalbard based on the presence of crevasse squeeze ridges (CSRs; also termed crevasse fill ridges; Sharp, 1985a) visible in glacier forelands. We strive to (i) identify which glaciers are surge- type glaciers and subsequently not optimal for traditional climate reconstructions and (ii) better understand the dynamics of Svalbard glaciers. Crevasse squeeze ridges (Sharp, 1985a; Benn and Evans, 2010; Rea and Evans, 2011) have been identified as a landform characteristic of surge-type glaciers and constitute an important component of the surg- ing glacier land systemmodel (Evans and Rea, 1999, 2003; Brynjólfsson et al., 2012; Schomacker et al., 2014). A detailed analysis of recent (2008–2012), high-resolution imagery from TopoSvalbard, provided by the Norwegian Polar Institute, allowed for an efficient survey of CSRs throughout Svalbard forelands. 1.1. Surge-type glaciers, landforms, and CSRs A surge-type glacier cyclically exhibits major fluctuations in ve- locity and length over timescales that range from a few years to sev- eral decades or centuries (Benn and Evans, 2010). Initially this Geomorphology 264 (2016) 52–60 ⁎ Corresponding author at: P.O. Box 156, Department of Arctic Geology, University Centre in Svalbard, UNIS, N-9171 Longyearbyen, Norway. E-mail addresses: wesleyf@unis.no (W.R. Farnsworth), olafur.ingolfsson@unis.no (Ó. Ingólfsson), mretelle@bates.edu (M. Retelle), anders.schomacker@uit.no (A. Schomacker). http://dx.doi.org/10.1016/j.geomorph.2016.03.025 0169-555X/© 2016 Elsevier B.V. All rights reserved. Contents lists available at ScienceDirect Geomorphology j ourna l homepage: www.e lsev ie r .com/ locate /geomorph behavior was believed to be caused by oscillations in basal condi- tions rather than external forcing by climate (Meier and Post, 1969; Sharp, 1988). Potential triggers for surges have been suggested, in- cluding reorganization of the basal hydrological system (Kamb et al., 1985), switching of thermal regime (Fowler et al., 2001; Sevestre et al., 2015), as well as sediment deformation at the glacier bed (Clarke et al., 1984; Björnsson, 1998). These different controlling factors have since been placed under one unifying theory relating to the enthalpy balance of a glacier (Aschwanden et al., 2012; Sevestre and Benn, 2015). Enthalpy balance ties together the hydrologic, ther- mal regime and bed deformation themes by highlighting a simple budget of energy exchange at the surface and base of a glacier as well as the frictional heat related to ice flow. Observations of surge events are relatively rare in Svalbard because of the remoteness of many glaciers and the extended surge cycles char- acteristic of the cold, dry region (Dowdeswell et al., 1991; Lønne, 2016). Commonly, surge-type glaciers are identified by the presence of specific landforms characteristic of a surge (Benn and Evans, 2010). The most frequently used landform accepted as an indicative fingerprint of surge activity is a looped medial moraine (Copland et al., 2011; Paul, 2015). Additionally, other features have also been used to indicate surge behavior, but are only unique to surge activity within a specific context. For example, trimlines above the ice surface are formed by surge-type and nonsurge-type glaciers but they have been used as surge indicators where the bulge of mass transferred from the accumu- lation zone to the ablation area formed a trimline of ice and debris (Sund et al., 2009; Kristensen and Benn, 2012). A surge trimline has a much lower preservation potential compared to a traditional trimline owing to the relatively high ice content. Extensive crevassing, high and wide- spread across the glacier that is not normally present, is indicative of a surge-type glacier during or shortly after a surge event (Hagen et al., 1993). Although these specific features are unique to surge events, their preservation potential after the surge is poor, as the crevasses dis- appear, the trimlines deteriorate and the loopedmoraines get reworked and distorted. Given the ‘switch on/switch off’ behavior of surge-type glaciers (Dowdeswell et al., 1995; Hansen, 2003), we feel these diagnos- tic landforms are sufficient evidence to identify surge-type behavior. Another landform unique to surge behavior that has not yet been widely used diagnostically for identifying a surge-type glacier is a cre- vasse squeeze ridge (CSR; Fig. 1). The CSRs were described in Iceland as elongated ridges ca. 1 to 2m in height and up to several hundredme- ters long (Sharp, 1985a,b) and were suggested to correspond with surge-type glaciers. The ridges, like looped moraines and trimlines, form during surge events but become exposed in surge-type glacier forelands during the quiescent phase. This delayed exposure allows for CSRs to potentially suggest surge behavior of a glacier decades to over a century after a single surge event (Fig. 1A). The ridges consist of matrix-supported diamict sediments and often form in networks on fluted till plains in forelands and can formoblique, transverse, or normal to ice flow (Sharp, 1985a,b; Kjær et al., 2008; Schomacker et al., 2014). The formation of CSRs is still not well understood (Rea and Evans, 2011; Ingólfsson et al., 2016). The relationship between the CSR and Fig. 1. (A)Oblique aerial image of Coraholmen, taken in July 2004, view toNE. Crevasse squeeze ridge network preserved subaerially on Coraholmen Islanddelimiting themaximumextent of the last Sefströmbreen surge ca. 1890 (Liestøl, 1993). For scale: Coastal cliffs are 8 to 10 m high. Box highlights prominent CSR network. Photograph: K.H. Kjær. (B) Image of CSRs (marked by arrows) melting out of the quiescent Von Postbreen, view to SE oblique to ice-flow, taken August 2012, by W.R. Farnsworth. Note three persons on the glacier in the background for scale, circled in red. (C) CSR preserved in the foreland of Kjerulfbreen marked with dotted lines, network visible in the background indicated with arrows, ridge in foreground roughly 0.5 m prominent composed of a matrix-supported diamict, view to the SE. Photograph: N. Aradóttir, August 2015. (D) Overview of CSR network (white box) preserved in the foreland of Kjerulfbreen, view to NE. Photograph: D. Ben-Yehoshua, August 2015. 53W.R. Farnsworth et al. / Geomorphology 264 (2016) 52–60 fluted till plain suggests several possible modes of formation (Sharp, 1985b; Bjarnadóttir, 2007; Ferguson et al., 2009). Similarmorphological structures can form from subaerial infill of moraine sediments (Morawski, 2005) or debris-rich thrust-faults (Glasser et al., 1998) but are not usually arranged in interconnected rhombohedral networks (Lovell et al., 2015b). Generally CSRs are believed to form by local satu- rated sediment infilling, from the bed upward, into basal crevasses dur- ing the latter part of a surge and subsequently meltout. In a detailed investigation of internal ice deformation structures, cross-cutting rela- tionships between compressional and extensional phases suggested a downglacier passage of a kinematic wave consistent with a surge (Lovell et al., 2015a). Rea and Evans (2011) concluded that high basal water pressure drives the infilling of basal crevasses from the bottom up. In theory high pore-water buildup correspondswith a fine sediment fractionation. In a detailed sedimentological study of the Elisebreen surge foreland where CSRs were identified, Christoffersen et al. (2005), describe a coarse till with a silt/clay composition ca. 35% com- pared to 65% sand and gravel. 2. Study area, climate, and glaciers Svalbard is an archipelago located along the dominant corridor of at- mosphericmoisture between theAtlantic and theArctic Basin, spanning from 74° to 81° N (Fig. 2A; D'Andrea et al., 2012). The islands are posi- tioned at the northern extent of the Gulf Stream (North Atlantic Drift) and the southern border of the Arctic sea ice front (Rogers et al., 2005). Presently, glaciers cover roughly 60% of the archipelago (Hagen et al., 2003), and Svalbard is categorized as having a dry High Arctic cli- mate with periglacial conditions, low winter temperatures, and warm continuous permafrost (French, 2007; Christiansen et al., 2010). Despite the archipelago's northern location, western Spitsbergen experiences a mild climate for its latitude where the relatively warm Norwegian Cur- rent controls weather and sea ice (Førland et al., 1997). Climate is char- acterized by the interactions between the Icelandic Low and Siberian High pressure systems, and as a result relatively high temperatures are driven north over Svalbard by the main North Atlantic cyclone track (Hanssen-Bauer et al., 1990). Svalbard precipitation is closely coupled to the mode of the North Atlantic Oscillation (NAO; Dickson et al., 2000) and falls predominantly in solid form. Longyearbyen, the main settlement (central Spitsbergen; Fig. 2) has a mean annual air temperature ca. −5 °C, and the annual sum of precipitation averages around 195 mm w.e. (Norwegian Meteorological Institute; http://www.eklima.no, n.d). Svalbard has over 2100 glaciers (Liestøl, 1993). These glaciers range in size from small cirque glaciers to large ice caps and exhibit a range of thermal properties (polythermal/cold/warm-based) aswell as terrestri- al and tidewater termini (Fig. 2B). The maximum Holocene extent of Svalbard glaciers and ice caps is generally considered to have occurred toward the end of the Little Ice Age (LIA; Werner, 1993; Humlum et al., 2003). The actual timing of the end of the LIA is dependent on whether it is defined by a warming trend in temperatures or the retreat of glacier fronts. The LIA on Svalbard culminated later (ca. 1930 CE) in comparison to Scandinavia, central Europe, and other parts of the world (Paasche and Bakke, 2010). The late culmination is visible in his- toric 1936 Norwegian Polar Institute oblique aerial photographs displayingmost icemargins up against or proximal to their LIAmoraine systems (Lyså and Lønne, 2001; Sletten et al., 2001). Since the end of the LIA retreat and thinning of Svalbard glaciers have been extensive through the twentieth century exposing an extensive recently glaciated landscape (Nuth et al., 2007). Surge-type glaciers exist regionally in dense clusters (Jiskoot et al., 2000; Grant et al., 2009), and Svalbard is one of these regions (Benn and Evans, 2010). Despite Svalbard clearly being recognized as a ‘surge hot spot’, how many of the glaciers on the archipelago actually exhibit surge behavior is not clear. Published estimates of the frequency of surge-type glaciers in Svalbard range from 90% (Lefauconnier and Hagen, 1991) to as low as 13% of all glaciers (Jiskoot et al., 1998). In our studywe highlight 277 documented surge-type glaciers based on previous inventories (Fig. 2B; Dowdeswell et al., 1991; Lefauconnier and Hagen, 1991; Hamilton, 1992; Hagen et al., 1993; Liestøl, 1993; Dowdeswell et al., 1999; Jiskoot et al., 2000; Dowdeswell and Benham, 2003; Błaszczyk et al., 2009; Sund et al., 2009). In contrast to previous studies, we choose not to address the percentages of surge-type glaciers in Svalbard or the subregions as we believe the values should corre- spondwith a temporal constraint. Given the dynamic ‘switch on/switch off’ of surge-type glaciers, we strictly present raw values of the number of glaciers where we have identified surge-diagnostic landforms. We have divided Spitsbergen, Nordaustlandet, and Barentsøya- Edgeøya into six regions, based on general location and climate (Fig. 2B). The distribution of documented surge glaciers is displayed in a pie chart in Fig. 2. A clear dominance of surge-type glaciers is visible in the southern part of Spitsbergen with about 45% of the documented surge-type glaciers located in this area. The other five regions account for the remaining ~55%. This apparent distribution could be the result of (i) relatively a higher frequency of observations of glacier front fluctuations over the last ~150 years through the southern region, or (ii) the climate/enthalpy balance in these areas might be more or less conducive to surge behavior (Sevestre and Benn, 2015). Northeast Spitsbergen, NW Spitsbergen, central Spitsbergen, Nordaustlandet, and the combination of Barentsøya/Edgeøya all have a similar number of documented surges despite these regions having a large range in the actual number of glaciers (Fig. 2B). 3. Methods Analysis of aerial imagery was carried out using TopoSvalbard (TopoSvalbard, 2016; http://toposvalbard.npolar.no/) provided by the Norwegian Polar Institute. The most recent aerial images available for Svalbard glacier forelands exist for years 2008, 2009, 2010, 2011, and 2012 depending on site and region. Aerial images allow detection of de- tails with up to 0.5-m resolution. The CSRs were interpreted based on relief, morphology, location, and cross-cutting relations within the gla- cier foreland system (Sharp, 1985b; Evans and Rea, 1999, 2003). These interpretations were in some cases compared to structures and the presence of CSR at known glacier forelands where the landforms have been studied in the field. The CSR networks were recorded and tabularized for glaciers on Svalbard for previously undocumented surge-type glaciers and for glaciers with documented surge histories (Dowdeswell et al., 1991; Lefauconnier and Hagen, 1991; Hamilton, 1992; Hagen et al., 1993; Liestøl, 1993; Dowdeswell et al., 1999; Jiskoot et al., 2000; Dowdeswell and Benham, 2003; Błaszczyk et al., 2009; Sund et al., 2009). Images from TopoSvalbard were analyzed for each foreland (Fig. 3). In the case where a glacier foreland had CSR but was unnamed, the glacier was coded by the region and a number (Fig. 3B). The unnamed glaciers were systematically Fig. 2. (A) Overviewmap of Svalbard located between the Arctic and Atlantic basins, base mapmodified fromwww.maps-world.net. (B) Themain islands of Svalbard divided into six general geographic regions. Blackdots indicate glacierswhere surge behavior has beendocumented (Dowdeswell et al., 1991; Lefauconnier andHagen, 1991;Hamilton, 1992;Hagenet al., 1993; Liestøl, 1993; Dowdeswell et al., 1999; Jiskoot et al., 2000; Dowdeswell and Benham, 2003; Błaszczyk et al., 2009; Sund et al., 2009). Inset pie chart displays the distribution of documented surge behavior by region. Base map modified from TopoSvalbard provided by the Norwegian Polar Institute. The three upper black dots on Vestfonna, Nordaustlandet are slightly offset on the map and correspond to glaciers to the north of map view: Søre Franklinbreen, Nordre Franklinbreen and Rijpbreen (from west to east). 54 W.R. Farnsworth et al. / Geomorphology 264 (2016) 52–60 numbered with the first glacier of each region being located farthest north while the final number corresponding to an unnamed glacier located furthest south in each region. For example, NWS-1 corresponds to an unnamed glacier with CSR in Northwestern Spits- bergen that is located in the upper part of the region. These glaciers are presented in the supplementary data (Table 2). 55W.R. Farnsworth et al. / Geomorphology 264 (2016) 52–60 4. Results Using aerial images from 2008 to 2012, CSRs have been identified in the forelands of 431 glaciers that have previously not been docu- mented to exhibit surge behavior. The CSRs have been identified in the front of undocumented surge-type glaciers in each of the six subregions of Svalbard: 164 in NW Spitsbergen, 55 in NE Spitsber- gen, 2 in Nordaustlandet, 40 in central Spitsbergen, 42 in the Barentsøya/Edgeøya region, and 128 in south Spitsbergen. Of these six regions the range in the number of glaciers with CSR is wide. This variability is likely caused by numerous factors including not only the amount of glaciated terrain but also the number of active ice margins. The crevasse squeeze ridges that have been identified in each of the 431 glacier forelands often are present in networks and commonly align transverse to ice flow but can in some cases form longitudinally as well (Fig. 3). The CSRs most often occur on basal till plains accompanied by, but perpendicular to, streamlined features like flutes, crag and tails, as well as the occasional eskers and drumlins (Fig. 3). Most of the CSR identified are located in the inner zone with some potentially transitioning into the intermediate zone of the surge-type glacier land system (Schomacker et al., 2014). At some glacier forelands the ridges were infrequent, melting out of the ice front or poorly preserved, mak- ing interpretation more difficult (Fig. 3B). The CSR are predominantly distributed around outlet glaciers where forelands have been exposed during the twentieth century ice retreat. This is in slight contrast to the documented (black dots) surge glaciers that in some cases are located at higher elevations, up on ice caps or ter- minating in fjords, without exposed forelands (Fig. 4). This variation in distribution is a result of the different methods used for identifying surge behavior. Crevasse squeeze ridges have been interpreted from fjord bathymetry (Flink et al., 2015; Lovell et al., 2015b), andmany tide- water glaciers have relatively long observation histories (Liestøl, 1993). Surging tributary glaciers at higher elevations have been mainly identi- fied by geodetic changes (Sund et al., 2009). In theory CSR can form and preserve at higher elevations, but the limiting factor seems to be the sediments at the glacier bed. Lønne (2016) discusses how variations in available sediment and deformable beds can aid or hinder the preserva- tion of sediment morphology. Lønne (2016) describes ‘Type B′ glaciers as land-terminating basins with poorly deformable substrate and low probability of proglacial moraine formation because of the coarse- grained debris. She notes that in these types of basins, the glacier debris rapidly transforms to alluvial or colluvial facies (Lønne, 2016). This type of glacier seems to correspond to higher alpine basins that do not form CSRs despite documented surge advances. The lowest value of forelands with CSR was identified in the Nordaustlandet region. Despite Nordaustlandet having the most exten- sive ice cover, it has relatively few active ice margins, especially com- pared to regions like NW and south Spitsbergen. A significant portion of the land-terminating ice caps on Nordaustlandet appear to have cold, passive margins owing to the cold dry climate and relatively low topography. Passive, cold-based margins are visible in aerial imagery of the region where widespread patterned ground and beach ridges are melting out from the nonerosive ice margins. The CSRs are present in front of active glacier margins, but under modern conditions very few land-terminatingmargins appear to be active. The exposure of pre- served beach ridges from the retreating ice margins may suggest that these terrestrial margins are not only passive today but may have been throughout the Holocene. Where the number of glaciers with CSR in most regions greatly exceeded the number of documented Fig. 3. Examples of CSR identified during the survey. (A) CSR network exposed in front of Keilhaubreen, South Spitsbergen. (B) CSR melting out of a small cirque in Firkantdalen, glacier coded SS-10. (C) Foreland of Kolfjellbreen, Central Spitsbergen, with CSR oblique to streamlined flow direction. (D) CSR network in the foreland of Kjerulfbreen, NW Spitsbergen. Field images of the same CSRnetwork are presented in (Fig. 1D),markedbywhite box. Aerial images of glacier forelandsmodified fromTopoSvalbard providedby theNorwegian Polar Institute. 56 W.R. Farnsworth et al. / Geomorphology 264 (2016) 52–60 surge-type glaciers, central Spitsbergen (likeNordaustlandet) displayed a relatively low number of forelands with CSR, compared to the number of documented surge glaciers (Fig. 4; Table 1). This relatively low num- ber is most likely a result of glacier front observations beingmuchmore common around the settlements of Spitsbergen and if a glacier front were to advance rapidly it would be less likely to go unnoticed. A great number of forelands with CSR were identified in NW Spits- bergen and in south Spitsbergen with 164 and 128, respectively (Table 1). These high values may relate to the obvious abundance of ac- tive ice margins and retreating glacier fronts exposing fresh forelands. The high concentration could also suggest optimal conditions for CSR formation or preservation, like deformable substrate. The CSRs were identified in the forelands of roughly two-thirds of the previously documented surge-type glaciers. The remaining one- third of the documented surge glacier forelands may lack CSR owing to a number of factors. Although a glacier has been documented to surge, the glacier (i) could be a tributary to a larger glacier, (ii) may ter- minate subaqueously, (iii) may not have preserved the CSR, or (iv) the landforms may not have formed because of other controls. In two forelands where CSRswere not identified, poor visibility or lack of aerial imagery made analysis impossible (Andrinebreen and Elfenbeinbreen, respectively). Fig. 4.Map of Svalbard divided into six subregions. Black dots indicate documented surge-type glaciers introduced in Fig. 2B. Red stars mark surge-type glaciers identified by the presence of CSRs in their forelands. Base map modified from TopoSvalbard provided by the Norwegian Polar Institute. The three upper black dots on Vestfonna are slightly offset on the map and correspond to glaciers to the north of map view; Søre Franklinbreen, Nordre Franklinbreen, and Rijpbreen (from west to east). Table 1 Distribution of CSR by region compared to documented surge events. Regions NW Spits. NE Spits. Nordaust. Cent. Spits. B/E-øya S. Spits. Total Doc. surges 77 22 11 42 14 111 277 CSR 164 55 2 40 42 128 431 Total surges 241 77 13 82 56 239 708 57W.R. Farnsworth et al. / Geomorphology 264 (2016) 52–60 5. Discussion The CSRs and concertina eskers are thought to be unique signatures of surge-type glaciers (Sharp, 1985b; Benn and Evans, 2010; Rea and Evans, 2011). However, Evans et al. (2012) suggested that CSRs and concertina eskers identified in the foreland of Hørbyebreen, Svalbard, could have formed frommeltwater pulses in a nonsurging polythermal glacier, driven by the buildup and release meltwater reservoirs. They conclude that the mapped Hørbyebreen land system does not agree with a surge-type glacier land system (Evans and Rea, 2003) and conse- quently provides an alternative explanation for the observed CSR and concertina eskers. We do not agree with this interpretation of Evans et al. (2012). A looped medial moraine is visible in the Norwegian Polar Institute 1936 oblique aerial image of Hørbyebreen, suggesting that the glacier has exhibited surge-type behavior. Consequently, our interpretation of the CSR and concertina eskers as landforms derived from a surge event is a simpler explanation for their origin than the interpretations of Evans et al. (2012). The distribution of surveyed CSR reflects the formation and preser- vation of the landforms as well as indirectly the limits of the method. The remote sensing method only allows for the analysis of the exposed terrestrial landscape. Thus subaqueous forelands and tributary glacier systems are not effectively evaluated in thismethod as CSR could poten- tially exist but are not visible in aerial imagery. Some glaciers may have documented surge behavior yet do not exhibit CSRs that are exposed terrestrially. In Nordaustlandet, 9 of the 11 documented surge-type gla- ciers (82%) do not have CSR (Table 2). These glacier systems predomi- nantly exhibit marine termini. The south Spitsbergen region also exhibits a large amount ofmarine-terminating glaciers, but the relative- ly high percentage of documented glaciers without CSRs (36%) is also a function of the minimal foreland exposure as many of the glaciers are tributary systems (Table 2). Another regionwith a high percentage of documented surge glaciers without CSR is in central Spitsbergen, where nearly 50% of the glaciers lack CSR (Table 2). Unlike Nordaustlandet and the south Spitsbergen regions, which are predominantly tributaries or marine terminating systems, these glaciers are small cirques that suggest that glacier size is a critical control in the formation of CSR. Brynjólfsson et al. (2012), suggested that CSRs are only identified on rare occasions in front of small cirque glaciers in alpine settings in Iceland. In the survey of glacier forelands in central Spitsbergen, numerous small glaciers exhibited rel- atively massive and extensive ice-cored moraines that currently corre- spond to small, near-empty cirques, consistent with the land system model of small surge-type cirque glaciers (Brynjólfsson et al., 2012). Despite glaciers needing to be of a minimum critical mass to form CSRs, the method is most effective with glaciers up to a certain size. The limiting factor of this method is exposed foreland which can be controlled by the size of a glacier, the length of a surge cycle, and the general sensitivity/response of the glacier system. For example, this method has been optimal for Svalbard because of the extended surge cycles characteristic of the region and the twentieth century warming driving significant retreat throughout the Arctic region (Dowdeswell et al., 1991; Nuth et al., 2007). The CSRmethodmay not be as successful in the Karakoram, where glaciers are significantly larger, surge cycles are shorter, and the debris-covered snouts are less sensitive to ablation (i.e. slow response time; Copland et al., 2011; Paul, 2015). Crevasse squeeze ridges are indirectly constrained by elevation as it seems deformable substrate (Lønne, 2016) and glacier sensitivity are limiting factors. Although it would be interesting to see if CSRs can be identified in regions like the Karakoram (given the importance of size, cyclicity, and sensitivity of surge-type glaciers), the CSRmethodwouldmost like- ly be more effective in other surge hotspot regions like Iceland, Alaska- Yukon, and Novaya Zemlya (Sevestre and Benn, 2015). The constraints and limiting factors on CSR formation still remain poorly understood. Numerous factors could influence the formation of CSRs, such as till grain-size distribution and effective pore-water pressures. Future studies should focus on whether the lack of CSR in the surge forelands is a function of preservation or CSR simply not forming because of topographic or basal sediment conditions. Focusing on potential characteristics that govern the distribution of CSRs (geolo- gy, bed conditions, and glacier size)would develop a better understand- ing of landform distribution and formation. Self-organizing maps (Kohonen, 2001) have been highlighted as a valuable technique for an- alyzing subtle relationships between disparate data sets (Fraser and Dickson, 2007). Although traditionally developed for finance, an in- creasing number of geoscience studies are using self-organizing maps, for example, to explore the relationship between streamlined glacial bedforms and the geological setting of the features (Dowling, 2016). Once the CSRs are formed, their preservation potential is relatively poor. The foreland of a retreating glacier (especially one that has surged) is a chaotic and unstable environment, where ice-cored terrain, glaciofluvial outwash, and high pore-water pressure in saturated sedi- ments result in an environment that is susceptible to reworking and deg- radation (Schomacker and Kjaer, 2007, Schomacker and Kjær, 2008). The formation and origins of CSR make them susceptible to degradation, and as a landform they have a relatively low preservation potential. The land- forms need to survive the initial deglaciation phase, when the ridges are exposed and highly susceptible to meltwater. They are subsequently exposed to periglacial processes where glacial terrain become reworked by frost weathering, slope processes, and degrading dead-ice bodies (Humlum et al., 2003; French, 2007; Schomacker and Kjaer, 2007, Schomacker and Kjær, 2008; Brynjólfsson et al., 2012). Thus, like other surge landforms, CSRs have a low preservation potential, but the window of time they can be used as surge indicators is notably later in the surge cycle than looped moraines or timelines. Of all the CSR surveyed, the majority formed around or within the last century and the landforms mostly have been exposed within the last 50 years or less depending on the size and sensitivity of the glacier sys- tems. Accordingly, Ingólfsson et al. (2016) pointed out that absence of a surge-type glacier land system does not unconditionally mean that a given glacier has not surged in the past. 6. Conclusions - This study concludes that the CSR survey is a simple, yet effective method for remotely identifying probable surge behavior. The CSRs have been identified in the forelands of 431 glaciers on Svalbard that have not previously been identified as surge-type. - This contribution is suggested to be a minimum estimate of glaciers that exhibit surge behavior on Svalbard because the presence of the CSR suggests surge behavior, but the lack of CSR in a foreland does not mean nonsurge behavior. Additionally, of the documented surge-type glaciers, CSRs are only visible in ca. 60% of their forelands. - The study utilizes open access data and remote sensing techniques to further knowledge of the presence and distribution of surge be- havior in Svalbard. - Results confirm a high density, but still unknown quantity of glaciers that exhibit surge behavior on Svalbard, at least through the late Holocene. - Sedimentological studies of CSRs (including grain-size properties and fabric analysis) are necessary to highlight the detailed processes of their formation. Additional focus could be placed on cross-cutting relationships of landforms in forelands to decipher their genesis in the subglacial environment. Acknowledgements We would like to recognize three anonymous reviewers and the editing staff for providing constructive comments and improving the manuscript. The authors additionally thank Kurt H. Kjær, Markus 58 W.R. 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Farnsworth et al. / Geomorphology 264 (2016) 52–60 Chapter V Holocene glacial and climate history of Svalbard - status, perspectives and challenges Farnsworth, W.R., Ingólfsson, Ó., Alexanderson, H., Forwick, M., Noormets, R., Retelle, M., Schomacker, A. (in prep.): Holocene glacial and climate history of Svalbard - status, perspectives and challenges. The cold-based, alpine ice cap, Bassen, northern Adventdalen, central Spitsbergen. 1 Holocene glacial and climate history of Svalbard - status, perspectives and challenges Wesley R. Farnsworth 1, 2* Ólafur Ingólfsson 1, 3 Helena Alexanderson 4, Matthias Forwick 2 Riko Noormets 1, Michael Retelle 5 Anders Schomacker 2 1Department of Arctic Geology, University Centre in Svalbard, UNIS Longyearbyen Norway 2Department of Geoscience, UiT, The Arctic University of Norway, Tromsø Norway 3Institute of Earth Sciences, University of Iceland, Askja, Sturlugata 7, IS-101 Reykjavík, Iceland 4Department of Geology, Lund University, Sölvegatan 12, Lund 223 62, Sweden 5Department of Geology, Bates College, Lewiston Maine USA *Corresponding Author: Wesley R. Farnsworth WesleyFunis.no KEYWORDS: Holocene, Spitsbergen, Neoglacial, Little Ice Age, glaciers, climate. Manuscript remains in preparation until database is finalized and ArcGIS supplement is completed. ABSTRACT: This study is a review of published literature of Holocene glacier and climate history from Svalbard and its surrounding waters. A database of (radiocarbon, terrestrial cosmogenic nuclide and optically stimulated luminescence) ages has been constructed, quality assessed and categorized by marine, terrestrial and lacustrine archive. Our review summarizes our understanding of glaciers and climate on Svalbard from the end of the Late Pleistocene (12.0 ka BP) to roughly the end of the Little Ice Age (LIA) as well as addresses gaps in our current knowledge. We (1) present a brief summary of major shifts in climate and ice cover across the Svalbard region throughout the Holocene; (2) introduce a quality assessed database of published ages that constrain glacier fluctuations (deglaciation, ice free, re-advance and marginal position) and climatic characteristics (warming ocean and ice cover expansion); (3) discuss challenges in methodology as well as potentials regarding sedimentary archives and finally (4) address the complexities of glacier systems and their dynamics in response to changes in climate. Furthermore, we identify some of current unknowns and propose possible prospects in order to approach these challenges in future studies. 2 INTRODUCTION Syntheses of accumulated field and geochronology data are prerequisite for putting new observations in context as well as re-assessing existing interpretations (Hughes et al. 2016). Empirical constraints are critical for developing and calibrating models that simulate palaeoclimate as well as past glacier dynamics and processes (Patton et al. 2017). Within this context, we synthesize published literature of Holocene glacier and climate history from the Svalbard archipelago and its surrounding waters. We introduce a database of empirical data specific to the Svalbard region as well as summarize the prominent climatic shifts and glacier fluctuations through the last 12,000 years (Fig. 1). The onset of the Holocene period (at 11.7 ka BP) marks a rapid transition from the end of the Pleistocene, characterized by a relatively cold period known as the Younger Dryas (YD), to the subsequent warmer interglacial conditions (Dansgaard et al., 1989; Wanner et al., 2008 Cohen et al. 2018). The global climate during the Holocene has traditionally been regarded as relatively stable, compared to the preceding Late Pleistocene (Dansgaard et al. 1993; Steffensen et al. 2008; Rockström et al. 2009), but this paradigm of a relatively uneventful Holocene is being increasingly challenged (Bond et al. 2001; Mayewski et al. 2004; Wanner et al. 2011). Reviews of Holocene climate and glacier oscillations have been synthesized globally (e.g. Solomina et al. 2015) as well as compiled for numerous Arctic and Alpine regions including the European Alps, Iceland, Arctic Canada and Greenland (Ivy-Ochs et al. 2009; Guðmundsson 1997; Geirsdóttir et al. 2009; Briner et al. 2016). In the north Atlantic, the Holocene climate has displayed a range of variability and fluctuations between warm and cold as well as humid and dry conditions on the multidecadal to multicentennial timescale (Mayewski et al. 2004; Wanner et al. 2011). A phrase coined ‘polar amplification’ refers to the sensitivity of Arctic regions to these climatic fluctuations and how these areas are greatly affected by small shifts in temperature and precipitation (Masso-Delmotte et al. 2013). By studying the rate and magnitude of past changes in climate we are able to better understand current climate dynamics, as well as more effectively predict antecedent climate scenarios (McKay & Kaufman 2014). This review summarizes the state of the art of glaciers and climate on Svalbard from the end of the Pleistocene to shortly after the end of the Little Ice Age (LIA) as well as addresses gaps in our current knowledge. While the end of the end of the Pleistocene and the Holocene with its three sub-divisions are defined (Greenlandian = Early, Northgrippian = Mid and Meghalayan = Late), the 3 review targets the time interval up until the early 20th century when a temperature record began in central Spitsbergen (Cohen et al. 2013; updated; Christiansen et al. 2013). Fig. 1 Location maps with A) inset map of the North Atlantic region with the Svalbard archipelago framed in black box. B) Overview map of Svalbard identifying key islands and regions. The warm West Spitsbergen Current (wsc) runs up along Svalbard’s western margin while the cool East 4 Spitsbergen Current (esc) traces down and around from the east. Figure maps modified from IBCAO and TopoSvalbard, respectively. We reconstruct Svalbard glacier and climate history by synthesizing geochronological data presented in published manuscripts detailing Holocene sedimentary archives (marine, terrestrial and lacustrine) and landforms. Our synthesis leads us up to the earliest historical ice front observations from Svalbard, and puts this accumulated Holocene data and observations in context with the ever- developing understanding of ice dynamics and the Arctic climate system. Thus, we strive to not only distinguish spatial trends in ice expansion of Svalbard glaciers and ice caps through the Holocene, but to correlate phases of ice-expansion with climatic periods, internal glacial processes, sea-level fluctuations, or some form of combination. Understanding how Arctic climate has fluctuated through the Holocene can provide insight as to how future changes may influence glaciers, sea level and potential environmental conditions. Svalbard has a long and rich history of scientific observations. The first field observations of Svalbard coincide with the earliest undisputed discovery of “Spitsbergen” by Willem Barentsz in 1596 (Clasezoon 1598; Hacquebord 1995; Arlov 2005). Although early settlers were initially drawn north to the coasts of Svalbard for hunting, whaling and trapping, by the mid-1800s scientific expeditions began accumulating detailed observations of the landscape, climate and ice-cover around the archipelago (Lottin et al. 1842; Bertrand 1852). In the mid-19th century numerous international scientific expeditions visited Svalbard eventually culminating in the first International Polar Year in 1882-1883 (Ekholm et al. 1887; Fig. 2). Through the last c. 150 years, the Svalbard archipelago has become a natural laboratory for observing the landforms, processes and dynamics of glaciers as well as the terrain which ice masses have shaped through the Quaternary (Holmström 1865; Ingólfsson 2011; Ingólfsson & Landvik 2013). The bulk of this work has focused on constraining the growth and break-up of the marine-based Svalbard-Barents Sea ice sheet in time and place (SBSIS; Boulton 1979; Blake 1962; Österholm 1990; Landvik et al. 1998; Mangerud et al. 1998; Hormes et al. 2013; Ingólfsson & Landvik 2013; Hughes et al. 2016). These reviews have predominantly focused on the Last Glacial Maximum (LGM) ice configuration, timing and disintegration and for clarity have often excluded data younger than the last termination. In several cases, studies of postglacial sea level and subsequent isostatic uplift have been used to better understand LGM ice cover and previous centers of mass, which indirectly has summarized chronological shoreline development through the Holocene (e.g. Salvigsen et al. 1981; Forman et al. 1990; Bondevik et al. 1995; Forman et al. 2004). 5 Holocene studies have long targeted unknowns of the glacier and climate history of Svalbard. Our synthesis addresses the following outstanding research questions: (1) is there evidence of Younger Dryas cooling and glacier re-advances on Svalbard? (2) Did meltwater from the collapsing Laurentide Ice Sheet (LIS) influence climate on Svalbard during the 8.2 ka BP event? (3) When was the Holocene (glacial) minimum and how extensive was the ice-cover during this period? (4) When was ice-cover the most extensive during the Holocene period and was the Little Ice Age the climax of the Neoglacial during the Late Holocene? (5) Which sedimentary archives provide the best detail of Holocene glaciers and climate? (6) To what extent has ice dynamics and surge-type behavior influenced Holocene glacier fluctuations on Svalbard? Setting Located along the dominant corridor of atmospheric moisture between the Atlantic and the Arctic Basin, Svalbard spans from 74° – 81° N (Fig. 1; Drange et al. 2005). At present, glaciers and ice caps cover roughly 57 % of the archipelago (Nuth et al. 2013). The region has a sensitive climate due to its position at the northern extent of the North Atlantic Drift (West Spitsbergen Current; Fig. 1) and the southern border of the Arctic sea ice front (Rogers et al. 2005). Svalbard is categorized as having a dry High Arctic climate with periglacial conditions, extreme winter temperatures and warm continuous permafrost (French 2007; Christiansen et al. 2010). Fig. 2. Photograph taken in Trygghamna in 1908 by Oscar Halldin. Glaciers Protektorbreen and Harrietbreen with ice-margins calving into the bay beyond Swedish Expedition Ships. From (De Geer 1908). Despite its high northern latitude, Spitsbergen currently experiences a relatively mild climate where the warm West Spitsbergen Current travels off the western extent of Svalbard and influences 6 weather and sea ice (Førland et al. 1997). Regional climate is directed by the interactions between the Icelandic Low and Siberian High pressure systems and as a result high temperatures are driven north over Svalbard by the main North Atlantic cyclone track (Hanssen-Bauer et al. 1990). Svalbard precipitation is closely coupled to the mode of the North Atlantic Oscillation (NAO; Dickson et al. 2000) and falls predominantly in solid form. The interactions of these air masses along the western flank of the archipelago now commonly drive winter conditions with warmer and wetter climate than normally expected as such latitudes (Førland et al. 1997; Eckerstorfer & Christiansen 2011). Sedimentary archives, landforms and geological reconstructions Holocene reconstructions of glaciers and climate represent a mosaic of data developed through a suite of sedimentary archives and landforms present in both marine and terrestrial environments. The different stratigraphic archives used to reconstruct glacier and climate history include terrestrial geological sections, marine sediment cores and threshold-lake records (Fig. 3). Geophysical data such as ground penetrating radar or chirp sub-bottom acoustic records can be paired with sediment cores and terrestrial stratigraphy, to extrapolate across larger areas. Additionally, the relative (and absolute) age of landforms and their cross-cutting relationships identified in submarine and subaerial data (marine bathymetry and aerial imagery) are used to reconstruct past glacier extent and other environmental conditions like relative sea level (Fig. 3). Despite over ten studies from 12 different ice-cores taken across Svalbard, ice core stratigraphy is discussed, but is not the focus of this review given the relatively short and young cryostratigraphic record (less than 1 ka BP; Isaksson et al. 2005; Grinsted et al. 2006; Divine et al. 2011). Methods Compilation of Ages Holocene geochronology from Svalbard and the surrounding waters that provide chronological evidence constraining glacier cover and marginal fluctuations or insight into climatic conditions over the last 12.0 ka were compiled into the SVALHOLA database (Table S1). Given the past variations in correction and calibration for the geochronological methods, we extent the databased to include ages in which mid-points of the 68th percentile fall below 12.0 ka BP as well as mid- points that fall above 12.0 ka yet error margins indicate the potential the ages is younger than 12.0 ka BP. Compiled ages and their meta-data were mined from scientific manuscripts, books, doctoral theses, geological reports (NP and Årsrapport / annual report from Oslo) and maps (Table 1). Dates obtained from compilations are cited as well as original source. To assure quality and consistency, 7 ages from MSc. theses were excluded unless subsequently referenced in a peer-reviewed manuscript. Fig. 3. Examples of Holocene landforms and sedimentary archives from De Geerbukta, NE Spitsbergen which provide insight into the glacial and climate history of Svalbard. Archives include marine, terrestrial and lacustrine records. Moraines formed on land or at the margin of a tidewater glacier have been identified and correspond to glacier re-advances. Raised marine sediments fingerprint post-glacial coastlines around Svalbard in some cases contain datable drift material (i.e. molluscs, whalebones and driftwood). Aerial images from TopoSvalbard and fjord bathymetry modified from Streuff et al. (2017). The SVALHOLA database compiles previously published ages of radiocarbon (14C, 14C AMS, 14C Conv.), luminescence (optically and infrared stimulated, OSL and IRSL) and terrestrial cosmogenic nuclide (TCN10Be, TCN26Al, TCN36Cl). Given the scope of the review and the Holocene focus, dating methods like thermo-luminescence (TL), electron spin resonance (ESR), U series and 210Pb have been omitted from the database. The database attempts to have a census date of 1 July 2018. We expect some published dates have been overlooked and hope any missing information can be updated and included in future versions of the SVALHOLA database. 8 Table 1. Presented metadata recorded for each date, included in the database (Table S1). Metadata from the database form the core of the criteria for quality assessment and palaeoglaciological classifications of each date as defined by the DATED1 database (Modified from Hughes et al. 2016). SVALHOLA ID -Unique database identification number Location -Country/sea, region, site name, SVALHOLA site number -Latitude and longitude co-ordinates: °N, °E (WGS84) -Comment on precision of location if not reported from original source. Sample characteristics -Site type: marine core, lake core, section, surface -Elevation (m a.s.l.) -Sample depth (m), if applicable Dated material -Sample field number and/or Laboratory ID number -Class of dated material: TPM (terrestrial plant macrofossils, including wood), organic (peat, detritus, bulk, mixed, aquatic macrofossils), bone (whalebone, tusks), shell (molluscs and mollusc fragments), foram (single species and mixed), sand, boulder, bedrock -Detailed description of dated material: free text -Organic material type: terrestrial (T), marine (M) Stratigraphic context -Detailed notes on stratigraphic setting: free text -Glacial context class: advance, margin, deglacial, ice free, exposure time (cumulative) Dating method -Radiocarbon (14C, 14C AMS, 14C Conv.), optically stimulated luminescence (OSL), infrared stimulated luminescence (IRSL), terrestrial cosmogenic nuclide (TCN10Be, TCN26Al, TCN36Cl) -Thermo-luminescence (TL), electron spin resonance (ESR), U series and Pb210 often fell out of the age range and scope of this review Quality control -Reliability of the age: 1 = reliable; 2 = possibly reliable; 3 = unlikely to be reliable (see Table 2 for criteria) Ages -Uncalibrated radiocarbon age / error (as reported, without correction for marine reservoir effect) -TCN age and error (as reported in source) -Calibrated/calendar age and error (reported to 1 SD). Radiocarbon ages calibrated to INTCAL13 or MARINE13 (Reimer et al. 2013) as appropriate (on basis of type of organic material: T/M). 10Be and 26 Al TCN ages recalculated using ‘Arctic’ production rate (Young et al. 2013) and Lal/Stone scaling (Lal 1991; Stone 2000). Necessary information to recalculate 10Be and 26Al TCN ages using different production rates additionally collated and recorded in Table S3 -Comments on calibration (e.g. beyond calibration curve limit) Comments -Any additional pertinent comments (e.g. reliability of date) Citation information -Source reference (author, year) -Compilation reference (author, year) -Database reference (for ages also included in other datasets, e.g. (Hormes et al. 2013; Mangerud and Svendsen 2017)) 9 Calibration of radiocarbon ages To present and evaluate the SVALHOLA dataset consistently, all radiocarbon dates were recalibrated with INTCAL13 and MARINE13 calibration curves using the Calib Rev. 7.0.4 program (Reimer et al. 2013). Calibrated age ranges are included for both the 68 and 95% probability in the database. In accordance with the DATED1 database we use the mid-point ± half of the total range at 68% probability to represent the calendar age and uncertainty for each age in our review (Hughes et al. 2016). The ages presented in text and figures are calibrated median ages in kilo-years before present (cal. ka BP) unless otherwise clarified. A marine reservoir age has been implemented for all marine samples by selecting ‘MARINE13’ and inputting a “Delta R” of 70 ± 30 (Table 1, S1; Mangerud & Svendsen 2017). Our chosen ΔR of 70±30 is specific for the Svalbard region (Mangerud et al. 2006; Mangerud & Svendsen 2017) and deviates from the DATED1 compilation, which for simplicity utilized a ΔR value of 0 for all marine samples from their reconstruction of the last Eurasian ice sheets (Hughes et al. 2016). Both radiocarbon ages and corrected ages are presented in the SVALHOLA data base (Table S1). Recalculation of terrestrial cosmogenic exposure ages Terrestrial cosmogenic exposure ages were (re-) calculated using the CRONUS-EARTH online calculator v3 (Balco et al. 2008; https://hess.ess.washington.edu) with the Arctic production rate calibration datasets (Young et al. 2013, 2014). An ‘Lm’ scaling (Lal 1991; Stone 2000) has been used in accordance with Youn et al. (2018). As argued by Hughes et al. and Young et al. (2016; 2018), no corrections have been made for post exposure uplift, erosion, or snow cover. Calculation of luminescence ages A luminescence age is calculated by determining the dose recorded by the grains and dividing it with the amount of dose received by the grains per unit of time (dose rate). The precision and accuracy of a luminescence age is dependent on a number of factors that affect these components and the largest uncertainties are typically related to geological factors rather than technical ones. For example, the dose rate is dependent on the amount of radioactive elements in the surrounding sediments, the water or ice content in the sediment as well as the depth below the ground surface (Rhodes 2011) and these factors must be measured and their changes through time estimated to arrive at an accurate age. The dose, on the other hand, can only be assumed to be correct if the resetting at the time of deposition was effective. If it was not, the grains are ‘incompletely bleached’ and some 10 luminescence signal from a prior event remains, making the measured dose and resulting age appear too high (old). This is not uncommon in depositional environments where sediment transport is short, deposition occurs rapidly or in turbid water or other setting where light is limited (Fuchs & Owen 2008). Experiments have shown that the TL signal bleaches more slowly than the IRSL signal which in turn is slower than the OSL signal (Godfrey-Smith et al. 1988; Alexanderson & Murray 2012a). In settings with limited light, it can therefore be argued that it is more likely that an OSL age gives a correct age than an IRSL or TL age. TL is nowadays considered less suitable for sediment dating and is not used much (Fuchs & Owen 2008; Wintle 2008); TL ages have therefore not been included in this compilation. For quality assessment criteria of OSL and IRSL ages, see Table 2. In the SVALHOLA database, ages are listed as presented in original publications (Table S2). Consistency and quality assessment of dates The SVALHOLA database consists of accumulated information acquired over the last 50 years. Dating techniques have developed over this period of time and the standard of what is considered “reliable” has risen (Hughes et al. 2016). We utilize the structure of the quality assessment criteria introduced in the DATED1 database (Table 2) to characterize the reliability of each age in the SVALHOLA database (Hughes et al. 2016). We have rated all ages depending on the dating technique on a 3-point system (quality 1-3) to rule-out potentially misleading ages. According to Hughes et al. (2016) a quality mark of; 1 = all criteria is met (likely reliable), 2 = some of the criteria are met but not all (probably reliable), 3 = no criteria are met (likely unreliable). Dates suggested being unreliable by original authors or other subsequent databases / compilations are rated quality 3. Dates of quality 1 and 2 standard have been used to reconstruct glacial and climate history while ages deemed quality 3 are presented in Table S1, but have been excluded from the developed reconstruction. We deviate from the DATED1 standard in one key aspect; sediment feeding marine molluscs (i.e. Portlandia arctica) are considered likely unreliable. Instead of receiving a lower rating, we strictly rate them as quality 3, thus excluding them from the reconstruction due to the high probability of an exaggerated age (England et al. 2013; Hughes et al. 2016). 11 Table 2. Age quality control criteria (based on Duller 2006, 2008; Thrasher et al. 2009; Wohlfarth 2009; Heyman et al. 2011; Alexanderson & Murray 2012b; England et al. 2013; Reimer et al. 2013). Ages within SVALHOLA are given a quality control (QC) rating based on the criteria specific to the dating method used. QC = 1, all criteria are satisfied; QC = 2, most of the criteria are satisfied; QC = 3 no (or few) criteria are satisfied (standard modified from Hughes et al. 2016). Dating technique Quality control criteria Radiocarbon 14C Conv (Conventional), 14C AMS -Known and uncontaminated sample material; sediment-feeding marine mollusc (e.g. Portlandia arctica) receives the lowest rating -Organic content >5% LOI -Sample composition: Conv - bulk samples not acceptable; AMS - bulk sample acceptable if age <20 ka -Within calibration range of INTCAL/MARINE13 -Uncalibrated 14C age determination provided with errors to enable recalibration using the latest calibration curves -Multiple and/or stratigraphically consistent ages Terrestrial cosmogenic nuclide TCN 10Be, 26Al, 36Cl -Multiple (ideally three or more, but at least two) samples from the same feature/site -Ages are internally consistent and clustered (reduced Chi-square value ~1) -Observed spread in ages is similar to expected measurement uncertainty -Geomorphological setting is accounted for: erosion, submergence, uplift -Data necessary to recalculate ages (10Be, 26Al) using different production rates (Balco et al. 2008) -No indication of isotopic inheritance, or if present expected/stated Luminescence OSL, IRSL -Quartz have a higher rating than feldspar-derived ages -Single-grain or small aliquot -Homogenous sample; preferably aeolian, fluvial, glacifluvial sediments that are likely to have received sufficient exposure. -Sample setting considered and accounted for; e.g. water-content history -Dose rate information and equivalent dose including errors described in source -Multiple and/or stratigraphically consistent ages All dating methods -Sample considered in situ, i.e. no post-depositional disturbance or reworking -Specified error margins -Precise ages: errors <10% of age -Details of geological and stratigraphical setting given -Considered by original authors to be reliable Palaeoglaciological interpretation and climatic association of dates We have simplified the palaeo-glaciolocical/climatological significance for a portion of the ages compiled in the SVALHOLA database based on metadata from the manuscripts or the author’s suggestions presented in the different studies. We assume all radiocarbon ages are sampled from ice free conditions. In some cases, additional classifications are placed on the ages. We choose not to classify palaeo-glaciological / climatological significance of individual ages taken from marine or lacustrine cores, but address their chronologies as a whole in both the results and discussion. Re-advance Radiocarbon ages of dateable material (i.e. vegetation, shell or bone) that has been reworked or overridden by an advancing glacier are classified as maximum constraining ages for a glacier re- 12 advance (Baranowski & Karlén 1976; Ronnert & Landvik 1995; Kristensen et al. 2009). In a study where numerous samples have been dated to constrain one glacier re-advance all ages are classified as a “re-advance” in the SVALHOLA database, but only the best constraining age (i.e. the youngest) is referred to in text and figures (Humlum et al. 2004; Table S1). Few Holocene glacier advances have been constrained in time by both maximum and minimum limiting ages (Lønne 2005; Farnsworth et al. 2018). Based on consideration by original authors ages classified as a “re- advance” are interpreted to correspond to a glacier fluctuation at or shortly following the mid-point of the youngest age (within the error margin). Deglacial Unlike previous reviews (Landvik et al. 1998; Ingólfsson & Landvik 2013; Hughes et al. 2016; Hogan et al. 2017) we choose not focus on the deglaciation in this paper. Furthermore, it is believed that most of the modern western and northern coastline of Spitsbergen had deglaciated by the onset of the Holocene (Hormes et al. 2013; Fig. 2a). In select locations we classify ages as “deglacial” where the site possesses stratigraphic information suggesting ice-free conditions shortly following ice-cover. Such dates include basal organic material from lake cores, the lowermost shells in marine cores from fjords and raised glacial marine sediments from fjord-heads (Alsos et al. 2015; Hald et al. 2004; Bartels et al. 2018; Farnsworth et al. 2018; Larsen et al. 2018). Additionally, the classification is used for radiocarbon ages of dateable material that has been sampled up-ice of modern margin (Blake 1989; Oerlemans et al. 2011) suggesting reduced ice-cover. Margin Although exposure age of erratic boulders is often used as an indication of the age of deglaciation (e.g. Hormes et al. 2011; 2013; Gjermundsen et al. 2013; Young et al. 2018), erratic boulder ages in the SVALHOLA database almost strictly target moraines and relate to an ice marginal position. Unlike ages association with a re-advance, where the youngest age is considered the best constraint, exposure ages corresponding to an ice margin often are presented in cumulative probability or in a histogram. The margin classification could also be used for a radiocarbon ages relating to an ice- contact delta. Climatic associations We assign associated climatic conditions to two types of radiocarbon samples, thermophile marine molluscs and terrestrial plants that have been entombed by (passive) ice (Mangerud and Svendsen 2017; Miller et al. 2017). We associate warm regional waters to thermophile marine molluscs 13 (TMM; Arctica islandica, Mytilus edulis, Modiolus modiolus and Zirfaea crispata) that have been sampled around Svalbard and radiocarbon date to different periods in the Holocene. The occurrence of these shallow marine molluscs, which are “effectively” extinct today, suggest warmer then present conditions and are indicators of the (marine) Holocene Thermal Maximum on Svalbard (Feyling-Hanssen 1955; Salvigsen et al. 1992; Hjort et al. 1995; Salvigsen 2002; Hansen et al. 2011; Farnsworth et al. 2017; Mangerud & Svendsen 2017). We classify radiocarbon ages of terrestrial plants that have been entombed by (passive) ice as indicators of ice-cover expansion (ICE; Miller et al. 2017). These samples contrary to ages associated to a “re-advance” do not correspond to unique/specific glacier systems. Large populations of these ages sampled over a vertical range have been associated with regional snow- line lowering, however a sole age does not indicate a conclusive driving factor of ice-cover expansion (summer temperature, winter precipitation, wind direction or some combination). Results The SVALHOLA database SVALHOLA contains 1707 individual dates from over 1200 discrete locations compiled from over 230 published sources (Fig. 5, Table S1–3). The spatial distribution is uneven and ages are dominantly from the western coast and fjords of Spitsbergen. The lacustrine archive has the lowest density of ages (~10 %) followed by marine cores (~29 %). Samples dated from terrestrial archives make up the final population of (~61 %). Over 90% of the dates in the SVALHOLA database are from radiocarbon analysis and roughly 50% of those samples are marine organisms. In the database, there are a total of 99 ages from Luminenscence dating (~6 %) while there is a total of 50 ages (~3 %) from TCN dating. Holocene dates from Svalbard are skewed in age toward the earlier half of the interglacial period. Roughly 60% of the SVALHOLA database dates between 12.0 – 6.0 ka BP with nearly half of the ages falling between 11.5 – 9.5 ka BP. Roughly 20% of the ages date to within the last 2 ka BP. Although the Mid Holocene has comparatively fewer dates per 500-year period, there is a relatively consistent distribution of ages, with a gentle decline from 9 – 2 ka BP. There is a lower percentage of “likely reliable” (quality 1) verse less reliable ages in the Early Holocene (12.0 – 8.0 ka BP). Holocene climate and glaciers Late Pleistocene – Early Holocene 14 Reworked dateable material of Late Pleistocene and Early Holocene age has been identified within and sampled from glacial deposits across Svalbard. Most often, this material is shell fragments re- sedimented in glacial (marine) diamict suggesting glacier override and in some cases glaciotectonism (Lønne 2005; Farnsworth et al. 2018). Several of the radiocarbon ages fall within the latter half of the Younger Dryas while the majority of ages date between 12.0 - 9.5 ka BP. The youngest age dates 9.1 ka BP. No clear clustering or synchronicity is visible in the age distribution of these reworked shells (Fig. 6). The oldest re-advances have been identified near the mouths of the fjords while younger Early Holocene re-advances are found in the inner tributaries and heads of fjords (Larsen et al. 2018). Fig. 4. Spatial distribution of ages compiled in the SVALHOLA database. Colored circles are 15 proportional to the number of dates of each method per site. Note the low density of chronological data from the eastern Svalbard and the east coast of Spitsbergen. The bulk of the ages are radiocarbon dates from low elevation coastal regions and in shallow waters to the west of Spitsbergen. Dates were compiled from citations present in Table S1 and at the end of this manuscript (full references are included in the supporting information; Data S4). Thermophilous marine molluscs that once inhabited Svalbard can be found preserved in raised marine sediments of Early Holocene age (Fig. 6; Feyling-Hanssen 1955; Salvigsen et al. 1992; Hjort et al. 1995; Salvigsen 2002; Blake et al. 2006; Hansen et al. 2011; Farnsworth et al. 2017; Mangerud & Svendsen 2017). 16 Fig. 5. Histogram of radiocarbon ages contained within the SVALHOLA database with dates presented with a 500-year bin size modified from 1000-year bin size presented by Hughes et al. (2016). Dates colored according to quality rating 1 (dark blue), 2 (light blue) and 3 (pale blue). Plot displays the same 1554 SVALHOLA dates with the mid-points stacked chronologically (black diamonds) and associated error bars (grey). Mid point ages extent over 12.0 ka BP as the database includes all ages in which error margins fall within the Holocene. The earliest radiocarbon ages of the warm water shell, Mytilus edulis date slightly before 11.0 ka BP and suggest the sea temperatures around Svalbard were roughly 2° C warmer than present (Fig. 6; Mangerud & Svendsen 2017). The occurrence of the Zirfaea crispata and Arctica islandica around Svalbard was used to compare modern conditions to the Holocene and suggests ocean temperatures peak between 10.0 – 9.2 ka BP and were at least 6° C warmer then present (Mangerud & Svendsen 2017). The Early Holocene terrestrial record is characterized by flights of raised marine beach sediments. These raised beaches indicate previous shorelines that have subsequently been uplifted relative to local sea level as a result of isostatic rebound (Forman et al. 1987; Bondevik et al. 1995). Early Holocene uplift rates have been derived from radiocarbon dated whalebones; drift wood and shells found on raised beaches (Salvigsen 1981; Häggblom 1982; Forman et al. 2004). Minimum rates derived from raised marine shorelines suggest 10 – 25 meters of uplift per 1.0 ka during the Early Holocene on Svalbard (Salvigsen 1981; Salvigsen & Österholm 1982; Forman et al. 2004). The occurrence of radiocarbon dated driftwood on the raised marine shorelines of Svalbard indicates an increasing rate of arrival between 12.0 – 10.5 ka BP which is followed by a stepped and variable decline in occurrence to 9.0 ka (Fig. 6). Peak driftwood arrival occurs between 10.5 – 11.0 ka BP (Fig. 6). Mid Holocene No (constraining) radiocarbon ages of re-sedimented material found within glacial deposits date to the Mid Holocene (Fig. 6). However, many thermophilous molluscs date to within the Mid Holocene. A slight decline in the occurrence of thermophilous marine molluscs suggests a short lived cooling following the marine Holocene thermal Optium (HTO) between 9 and 8.2 ka BP, however ocean temperatures still remained roughly 2° C warmer than present (Mangerud & Svendsen 2017). The persistence of the Mytilus edulis and the Modiolus modiolus in the Mid Holocene suggests Atlantic water continued to reach Svalbard and ocean temperatures were c. 4° C warmer than present between 8.0 – 6.5 ka BP (Salvigsen et al. 1992; Salvigsen 2002; Blake et al. 17 2006; Mangerud & Svendsen 2017). This period of stable warm ocean conditions is followed by a gradual decrease in temperatures until the end of the Mid Holocene where values reach those comparable to modern (Mangerud & Svendsen 2017). The presence of driftwood on raised Mid Holocene shorelines remains lower than the peak in occurrence from the Early Holocene, yet suggests consistent arrival of material from roughly 9.0 – 6.5 ka BP. This interval is followed by a decline in the arrival of driftwood at 6.5 ka BP and continuing through the end of the Mid Holocene (Fig. 6). Fig. 6. Plot of the 930 radiocarbon ages from terrestrial archvies (including raised marine) from the SVALHOLA database with the mid-points stacked chronologically (black diamonds) and 18 associated error bars (grey) modified from Fig. 5. The Younger Dryas period is marked with a blue column while the Early and Late Holocene is divided by a grey column delimiting the Mid Holocene. Specific radiocarbon ages from the database have been highlighted based on palaeo- glaciological / climatological associations. Dateable material reworked in glacial sediments (blue); thermophilous marine molluscs (red) and ice-entombed moss (green) are represented by enlarged colored circles with error whiskers. Histogram plot at base indicates the number (n=168) of radiocarbon dated driftwood samples from Svalbard presented within a 500 year bin size (modified from Dyke et al. 1997). The youngest bin (modern – 0.5 ka BP) has been excluded due to lack of studies focusing on driftwood from modern shorelines. Late Holocene There are an increasing number of radiocarbon dates from material re-sedimented in or overlain by glacial deposits throughout the Late Holocene (Fig. 6). Although some Late Holocene samples are of re-worked shells (Punning et al. 1976; Sharin et al. 2014; Farnsworth et al. 2017), roughly half of the samples are from over-ridden (in situ) vegetation (Baranowski & Karlén 1976; Dzierzek et al. 1990; Furrer 1991; Humlum et al. 2005). These samples have been collected across Svalbard and ages become more frequent in the last 2 millennia (Fig. 6). Over half of the ages date to between 1.0 – 0.5 ka BP (Fig. 6). In addition to material that has been reworked or overridden by glacier, plants that have been entombed in passive ice (cold, perennial snow patches and fonna / thin ice-caps) have also been sampled and dated (Fig. 6). Over 40 samples of moss and vegetation preserved under cold- based ice have been collected from retreating modern ice margins in central Spitsbergen (Miller et al. 2017). While the oldest ages range back to nearly 4 ka BP, c. 80% of the samples are younger than 2.0 ka BP (Fig. 6). The occurrence of thermophilous molluscs decreases through the Mid Holocene and tapers off entirely at the start of the Late Holocene (Fig. 6). In addition to the final two Mytilus edulis dating c. 3.7 ka BP, it appears the species may have returned to inner Isfjorden based on a single young sample dating 0.9 ka BP (Fig. 10; Samtleben 1985; Mangerud & Svendsen 2017). During the first half of the Late Holocene driftwood occurrence remains low (Fig. 6). The period between 2.5 - 2.0 ka BP marks the Holocene minimum in driftwood arrival and matches the occurrence rate seen during the end of the Younger Dryas (Fig. 6). A sharp increase in driftwood arrival is exhibited during the final two millennia. The driftwood count from the most recent 500- years (bin 0.5 ka BP to modern) has been excluded given the general lack of studies targeting radiocarbon dating driftage from the modern shorelines. 19 Holocene sediment chronologies from lakes and fjords On Svalbard, the resolution of lake and fjord core chronology ranges widely owing to core length and datable material (Fig. 7). Based on a collection of the highest resolution lake and fjord chronologies from published Svalbard Holocene studies, lake cores generally tend to be shorter in length (avg. 2.4 m) than fjord cores (avg. 4.9 m) and average a greater amount of dates (Fig. 7; 4.6 verse 1.8 dates per meter, respectively). Fig. 7. Diagram highlighting the age and distribution of dates used to constrain the highest resolution lake (n=11) and fjord (n=8) sediment cores used for Holocene environmental reconstructions around Svalbard. Bulk sediment (white boxes), terrestrial plant macrofossils (black boxes), crypto-tephra (green boxes), marine molluscs and foraminifera (blue boxes) are used to constrain core chronologies. Values beneath chronologies indicate the number of dates per core length of each study. 20 Holocene environmental reconstructions from marine sediments often use mollusc shells or foraminifera for radiocarbon dating and constraining core chronologies. In lake sediment cores, a mixture of material has been used to constrain chronologies (Fig. 7). Early studies relied on radiocarbon dating bulk sediment (e.g. Hyvärienen 1970), while modern studies target terrestrial plant macrofossils (de Wet et al. 2017) and occasional mollusc shells sampled just below the lacustrine-marine boundary in isolation basins (Svendsen & Mangerud 1997). Recent studies have used tephra and crypto-tephra to constrain lake chronologies where terrestrial macrofossils have been sparse (Fig. 7; D`Andrea et al. 2012; van der Bilt et al. 2017). DISCUSSION The three Holocene sub-divisions; Early, Mid and Late Holocene, presented in this review correspond to the Greenlandian, Northgrippian and Meghalayan respectively (Cohen et al. 2013; updated). For each Holocene stage, we discuss findings from the marine, terrestrial and lacustrine archives respectively. The stage intervals are from 11.7 – 8.2 ka BP, 8.2 – 4.2 ka BP and 4.2 ka BP to present. We introduce these three sections by briefly discussing the conditions at the transition, from the end of the Pleistocene into the Early Holocene. We close this summary describing the transition from the end of the Little Ice Age into the 20th century as this marks the start of the rich observational record of Svalbard climate, ice cover and landscape. Despite certain micro-climatic variability, we choose not to sub-group data, but generally discuss ages for the entire Svalbard region. In some cases, we refer to locations at the mouth of a fjord relative to the tributary of head of the fjord-system. Late Pleistocene - Holocene transition Maximum ice cover during the last 12 ka was likely around 12 ka BP, as residual SBSIS still covered a large portion of Svalbard (Fig. 8). The exact ice extent is not well understood given the sparse ice marginal positions known from this period. Additionally, it is unknown how tightly the deglaciation ages from marine sediment cores and raised marine sediments constrain the actual deglaciation as these ages are minimum constraining values and could potentially reflect conditions centuries after the actual deglaciation (Larsen et al. 2018). Several recent studies present new radiocarbon ages that suggest less extensive ice cover at 12.0 ka BP (Farnsworth et al. 2018; Larsen et al. 2018) than the DATED-1 time slice suggests (Fig. 8; Hughes et al. 2016). Furthermore, the resolution of these ice extent reconstructions does not effectively depict the detail of residual ice-cover during the transition from Pleistocene to Holocene. For example, no 21 terrestrial evidence from the Seven Islands, northern Svalbard, suggests ice free conditions prior to the Early Holocene (polygon with dashed line; Fig. 7). The earliest deglaciation ages are from raised beaches dating to c. 11.0 ka BP and at present the Seven Islands are not glaciated (Forman & Ingólfsson 2000). Fig. 8. IBCAO map of Svalbard with polygons reflecting ice cover at different time-slice reconstructions during the Lateglacial and Early Holocene. The most-credible time-slice reconstructions from Hughes et al. (2016) are from the DATED-1 database of Eurasian ice sheets, while the marginal positions from Mangerud & Landvik (2007) and Hormes et al. (2013) are modified from studies. Red circles mark sample locations where recent radiocarbon ages suggest ice free conditions prior to 12.0 ka BP (Farnsworth et al. 2018; Larsen et al. 2018). Red shaded polygon over the Seven Islands highlights reconstructed ice cover that existed until the onset of the Holocene (but lacks from other reconstructions due to resolution). 22 Younger Dryas-The period between 12.9 and 11.7 ka BP is believed to have been a phase of pronounced cooling interpreted from a marked boundary in Scandinavian pollen records, indicating the return of the Dryas octopetala, a cold-tolerant plant (Andersson 1896; Rasmussen et al. 2006) Interpretations of δ18O from Greenland ice core records suggest a rapid ~10°C deviation in atmospheric temperature during this time interval (Carlsen 2013). Data suggest that this period was not only characterized by pronounced cooling, but the event was a rapid climate shift and may have had implications for larger regions than just the North Atlantic (Dansgaard et al. 1993; Isarin & Renssen 1999). Studies have gone on to search for signatures of the Younger Dryas globally, suggesting contemporaneous (and synchronous) glacier re-advances in the European Alps, the North American Rockies as well as shifts in lake chemistry records from Chile to lake Baikal are believed to evidence a global response to this rapid climate event (Bennet et al. 2000; Ivy-Ochs et al. 2009; Davies et al. 2009). Although the North Atlantic is interpreted to have been the epicenter of this rapid climate event and despite the far-reaching fingerprints of the Younger Dryas cooling suggested by some studies, no conclusive evidence of any such event has been identified on Svalbard (Hormes et al. 2013). Similarly, our review finds no conclusive evidence for a Younger Dryas cooling driving synchronous glacier re-advances during any period between 12.9 – 11.7 ka BP. Radiocarbon ages from this time interval predominantly correspond to shells found in raised marine sediments suggesting ice-free conditions and active beach formation (Landvik et al. 1987; Forman et al. 1990; Eitel et al. 2002; Bruckner et al. 2003; Landvik & Mangerud 2007). Recent studies have also suggested substantial ice loss constrained to the Younger Dryas time interval thus, instead of glaciers re-advances or marked standstills, evidence in some locations suggests the contrary (Fig. 8; Hogan et al. 2017; Farnsworth et al. 2018; Larsen et al. 2018). Several radiocarbon ages, often of shell fragments re-sedimented in glacial (marine) sediments suggest that glacier re-advances occurred during the latter half of the Younger Dryas period (Fig. 6). These glacier re-advances are generally not well constrained in time (lacking both maximum and/or minimum age constraints) and exhibit no clear synchronicity (Fig. 6; Mangerud et al. 1992; Rasmussen & Thomsen 2014; Farnsworth et al. 2017, 2018; Larsen et al. 2018). Early Holocene (11.7 - 8.2 ka BP) By the beginning of the Holocene, the majority of the marine based sectors of the SBIS had already collapsed, and ice margins had retreated back inside the modern coastline of Svalbard (Mangerud et al. 1992; Hormes et al. 2013; Hughes et al. 2016). Deglaciation continued in a time transgressive 23 manner, characterized by initial retreat through the fjords and subsequently in the fjord-valleys and terrestrial realms (Ingólfsson & Landvik 2013; Gilbert et al. 2018). Additionally, during this period, summer insolation approaches its Holocene maximum in the northern hemisphere (Laskar et al. 2004). Marine- Data from the marine archives suggest early and marked warm regional conditions driven by the incursion of Atlantic waters around Svalbard (Hald et al. 2004; Mangerud & Svendsen 2017). Evidence from marine microfossils suggests warm regional condition prior to the onset and throughout the Early Holocene (Hald et al. 2004, 2007; Ślubowska-Woldengen et al. 2007; Skirbekk et al. 2010; Jernas et al. 2013; Łącka et al. 2015). Marine sediment core chronology from fjords in western and northern Svalbard suggest high sedimentation rates with regular outsized particles interpreted as ice rafted debris (IRD; Hald et al. 2004; Flink et al. 2017; Bartels et al. 2017; 2018). Reconstructions of sea ice cover during the onset of the Early Holocene are characterized by an abrupt decrease in sea ice proxy IP25 suggesting a decline in sea ice extent (Müller & Stein 2014; Bartels et al. 2018). Terrestrial- The Early Holocene exhibits the greatest rates of isostatic uplift recorded in the last 11.7 ka BP. Minimum rates derived from dated raised marine shorelines suggest 10 – 25 meters of uplift per 1.0 ka during the Early Holocene on Svalbard (Salvigsen 1981; Salvigsen & Österholm 1982; Forman et al. 2004). The occurrence of driftwood found on Arctic shorelines has been used as a proxy for semi- permanent sea ice concentrations (Häggblom 1982; Funder et al. 2011; Hole & Macias-Fauria 2017). Where sea ice cover is too low no driftwood arrives, while with multiyear sea ice, driftwood is shielded from the shorelines (Funder et al. 2011). The initial growth in driftwood arrival could be related to sea ice concentrations transitioning from permanent to semi-permanent cover at the onset of the Holocene as well as the increased availability of deglaciated shorelines as catchment for driftwood (Fig. 6; Häggblom 1982): The records of raised marine organisms identified in Early Holocene sediments across the coastal regions of Svalbard also corroborates with the marine data. Reviews of thermophilous marine molluscs found around Svalbard also indicate early and exceptionally warm regional waters (Salvigsen et al. 1992; Salvigsen 2002; Blake et al. 2006; Mangerud & Svendsen 2017). The occurrence of the Zirfaea crispata and Arctica islandica only existing in the Svalbard raised marine record between 10.0 - 9.2 ka BP suggests that the maximum ocean temperatures experienced in the 24 Holocene occurred in the beginning, slightly prior to peak summer insolation (Lasker et al. 2004; Mangerud & Svendsen 2017). Lacustrine- Few lacustrine chronologies on Svalbard begin prior to the onset of the Holocene (Fig. 6). Alkenone and hydrogen isotope reconstructions have been used as proxy records from lake sediments collected in northwestern Spitsbergen and suggest warm, moist conditions as early as 12.8 cal. ka BP (Balascio et al. 2018; van der Bilt et al. 2018; Gjerde et al. 2018). Most of the investigated lakes became ice-free (Holmgren et al. 2010; Alsos et al. 2015; Røthe et al. 2015; van der Bilt et al. 2016; de Wet et al. 2018) or isolated from the marine environment during the Early Holocene (Snyder et al. 1994; Svendsen & Mangerud 1997; Røthe et al. 2018). Lacustrine records from Svalbard often exhibit high mineragentic fractionation and sedimentation rates that taper off through the Early Holocene (Snyder et al. 1994; Svendsen and Mangerud 1997; Røthe et al. 2015; van der Bilt et al. 2016; de Wet et al. 2018; Røthe et al. 2018). Holocene marine thermal optimum and re-advancing glaciers Although the majority of the terrestrial landscape is still evacuating SBSIS-ice, the marine seas around Svalbard undergo an Early Holocene Thermal Optimum roughly between 10.0 and 9.2 ka BP (Hormes et al. 2013; Mangerud & Svendsen 2017). As the landscape is dominantly engulfed by ice, it is safe to say the marine and terrestrial environments are climatically out of phase at this period in the Holocene. Glacier response to Early Holocene conditions appears complex. In general, marine terminating glaciers responded to the warming by retreating back to land while cirque glaciers greatly diminish in size and in some cases disappeared completely (Svendsen & Mangerud 1997; Snyder et al. 2000; Forwick & Vorren 2009; de Wet et al. 2018). On Nordaustlandet, mollusc shell samples from within a thrusted debris band located 6 km inside of the Søre Franklinbreen ice margin date 10.3 ka BP and suggest that the NE outlet of Vestfonna had retreated to at least this position early in the Holocene (Blake 1989). Despite the progressive deglaciation and imminent retreat characterizing the Early Holocene, evidence from glaciers of varying size found across Svalbard, suggests asynchronous ice margin re-advance (Fig. 8; Salvigsen et al. 1990; Mangerud et al. 1992; Ronnert & Landvik 1993; Brückner et al. 2002; Eitel et al. 2002; Lønne 2005; Farnsworth et al. 2017, 2018; Larsen et al. 2018). Although the exact magnitudes of the re-advances are unknown, ice margin extent is often several kilometers outside of Late Holocene glacier maxima (Lønne 2005; Farnsworth et al. 2018). Similar to the late Younger Dryas, glacier re-advances are most often interpreted from shell 25 fragments re-sedimented in glacial (marine) sediments suggesting glacier override, reworking and in some cases glaciotectonism (Lønne 2005; Farnsworth et al. 2018). The oldest re-advances have been identified near the mouths of the fjords while younger Early Holocene re-advances are found in the inner tributaries and heads of fjords (Larsen et al. 2018). Generally, the timing of the re- advances seems to follow the time transgressive deglaciation (Landvik et al. 2014; Farnsworth et al. 2018). The final glacier re-advance identified in the first half of the Holocene deposited shells 180 m a.s.l. in inner Wijdefjorden (c. 6.5 km inside of the present margin; Klysz et al. 1988). Although the extent of the re-advance interpreted from these deposits is unknown, the data suggests not only had the palaeo-outlet glacier undergone significant Early Holocene retreat, it re-advanced to a marginal position un-matched at any point in the Late Holocene. Still today, there are numerous pre-Mid Holocene moraines that have been identified in marine and terrestrial environments, yet still are undated (Salvigsen & Österholm 1982; Forwick & Vorren 2010; Henriksen et al. 2014; Røthe et al. 2015 Flink et al. 2018; Farnsworth et al. 2018). Mid Holocene (8.2 – 4.2 ka BP) After summer insolation peaked during the Early Holocene, northern hemisphere insolation values begin a decreasing trend that progresses throughout the rest of the Holocene (Laskar et al. 2004). Marine- Data from the marine archive suggests a general decline in ocean temperatures during the Mid Holocene. Evidence from marine microfossils suggests cooling of regional waters during this period (Hald et al. 2004; 2007; Ślubowska-Woldengen et al. 2007; Skirbekk et al. 2010; Jernas et al. 2013; Łącka et al. 2015). Marine sediment cores exhibit a decline in sedimentation rate and IRD dissipates in most fjords (Forwick & Vorren 2009; Skirbekk et al. 2010; Kempf et al. 2013; Nielsen & Rasmussen 2018). Reconstructions of sea ice cover derived from IP25 records spanning the Mid Holocene suggest a progressive increase in ice-cover starting around 8.0 ka BP (Müller & Stein 2012). Terrestrial- Similar to the Early Holocene, most of our understanding of the Mid Holocene environment on Svalbard is from raised marine sediments and driftage. Mid Holocene raised marine shorelines suggest glacio-isostatic uplift rates decline throughout the Mid Holocene on the order of 5 m per 1 ka (Bondevik et al. 1995; Forman et al. 2004). In several locations on the northern and western coasts of Svalbard, Mid Holocene transgressions have been interpreted through a combination of shoreline morphology and chronology (Forman 1990; Forman & Ingólfsson 2000). 26 The presence of driftwood on raised Mid Holocene shorelines remains lower than the peak in occurrence from the Early Holocene (Fig. 6), yet suggests consistent arrival of material from roughly 9.0 – 6.5 ka BP. This interval is followed by a gentle decline in the arrival of driftwood towards the end of the Mid Holocene (Fig. 8). The pattern of driftwood occurrence through the Mid Holocene seems to loosely follow trend with declining ocean temperatures reconstructed from marine microfossils as well as inversely correlated with sea ice reconstructions derived from IP25 (Hald et al. 2004, 2007; Ślubowska-Woldengen et al. 2007; Skirbekk et al. 2010; Müller &Stein 2012). A slight decline the occurrence of thermophilous marine molluscs suggests a short lived cooling following the marine HTO between 9 and 8.2 ka BP, however ocean temperatures still remained roughly 2° C warmer than present (Mangerud & Svendsen 2017). The persistence of the Mytilus edulis and the Modiolus modiolus in the Mid Holocene suggests Atlantic water continued to reach Svalbard and ocean temperatures were c. 4° C warmer than present between 8.0 – 6.5 ka BP (Salvigsen et al. 1992; Salvigsen 2002; Blake et al. 2006; Mangerud & Svendsen 2017). This period of stable warm ocean conditions is followed by a gradual decrease in temperatures until the end of the Mid Holocene where values reach those comparable to modern (Mangerud & Svendsen 2017). Lacustrine- Lake sediment records spanning the Mid Holocene suggest high productivity and low sedimentation rates (Birks 1991; van der Bilt et al. 2015; Alsos et al. 2015; Gjerde et al. 2018). Chronologies from modern glacial lakes suggest greatly reduced or ice-free catchments during the Mid Holocene characterized by organic rich strata with minimal mineragenic accumulation (Svendsen & Mangerud 1997; Snyder et al. 2000; Røthe et al. 2015, 2018; de Wet et al. 2018). Glaciers retreat (entirely?) to the Holocene Glacial Minimum while waters cooled Based on glacial lake records and the summer insolation curve, we presume that Svalbard glaciers retreated back to their Holocene minimum during the early-Mid Holocene (Fig 7; Laskar et al. 2014). However, we have no exact spatial constraint on any Svalbard ice margins during the Mid Holocene. We assume if glaciers existed in the Mid Holocene, they were more reduced then any Late Holocene glacier extent on Svalbard. We therefore have no morphological evidence (frontal/marginal moraines) identifying any glacier re-advances during this period (Fig. 6). Two marine stratigraphy studies have suggested Mid Holocene glacier activity at the mouth of Isfjorden and western Nordaustlandet based on fluctuating rates of mineragenic sedimentation (Forwick & Vorren 2007; Kubischta et al. 2011). It is inconclusive if marine stratigraphy from these sites relates 27 to glacier re-advances or is a result of submarine mass failures or enhanced melting from proximal glacier systems. Glacial lake studies from the west coast of Spitsbergen suggest there was no glaciogenic input from their corresponding cirque glaciers during the Mid Holocene (Fig. 9; Linnébreen, Svendsen & Mangerud 1997; Linnébreen, Snyder et al. 2000; Kløsa, Røthe et al. 2015; Hajeren, van der Bilt et al. 2016; Gjøavatnet, de Wet et al. 2018). Terrestrial vegetation thrusted up into tidewater glacier ice suggests the marine terminus had retreated back from the fjord, at the mouth of Hornsund prior to 8.0 ka BP (Oerlemans et al. 2011). Fig. 9. Maps of Holocene minimum ice extent on Svalbard c. 8-7 ka BP. A) Glacial lakes records from Gjøavatnet (G), Hajeren (H), Kløsa (K) Vårfluesjøen (V) and Linnévatnet (L), suggest ice-free or greatly reduced glaciers in catchments during the Mid Holocene (Svendsen & Mangerud 1997; Snyder et al. 2000; Røthe et al. 2015; van der Bilt et al. 2016; Røthe et al. 2018; de Wet et al. 28 2018). Reduced ice rafted debris in marine sediment cores from Kongsfjorden (Kf), Van Kuelenfjorden (Vk), Woodfjorden (Wf) and Storfjorden (Sf) suggests tidewater glaciers have retreated back to terrestrial margins (Skirbekk et al. 2010; Kempf et al. 2013; Bartels et al. 2017; Nielsen & Rasmussen 2018). B) Ice rafted debris in marine sediment records suggest tidewater termini persistent in Billefjorden (JM97-943), Tempelfjorden (JM02-977), C) Van Mijenfjorden (MD99-2305) and D) Wahlenbergfjorden HH12-13 & GeoB10831-3) throughout the Holocene (Hald et al. 2004; Baetan et al. 2010; Forwick & Vorren 2010; Flink et al. 2017; Bartels et al. 2018). There are over 850 km of marine terminating glacier margin in Svalbard today, (Blaszczyk et al. 2009). However, it has been speculated if any tidewater glaciers survived the Mid Holocene as most marine sediment records from Svalbard glacial-fjords exhibit reduced or zero Mid Holocene IRD (Fig. 9; Isfjorden, Forwick & Vorren 2009; Kongsfjorden, Skirbekk et al. 2010; Van Kuelenfjorden, Kempf et al. 2013; Woodfjorden, Bartels et al. 2017; Storfjorden, Nielsen & Rasmussen 2018). However, ice rafted debris has been observed through the entire Holocene in sediment cores from several Svalbard fjords including Van Mijenfjorden, Billefjorden, Templefjorden and Wahlenbergfjorden (Hald et al. 2004; Forwick and Vorren 2009, 2010; Baeten et al. 2010; Flink et al. 2017; Bartels et al. 2018). Ice rafted debris is a strong indicator of marine terminating glaciers within a fjord-system but has also been associated with sea ice both as a raft for beach sediments, as well as a (semi-) permanent restraint limiting ice rafting (Forwick & Vorren 2009). If we assume this IRD relates to tidewater glacier systems, the IRD records indicate, Svalbard glaciers not only survive the entire Holocene interglacial, but several Svalbard glaciers never entirely retreated from their fjord-heads. A new modeling study based on Holocene relative sea level concludes that Svalbard glaciers survived the HTO and suggests Nordaustlandet and eastern Spitsbergen were the main regions that hosted this Mid Holocene ice (Fjeldskaar et al. 2018). These regions are characterized by cooler climate based on elevation and distance from the West Spitsbergen Current (Fig. 1; Førland et al. 2011). Although the Van Mijenfjorden region has summits reaching over 1000 m a.s.l. it generally does not reflect these same cool characteristics. While Paulabreen in southeastern Van Mijenfjorden is the primary source of IRD in the fjord at present; it was not necessarily the dominant or only source of IRD during the Early and Mid Holocene (Fig. 9). Based on the size of the Reindalen and Kjellströmdalen valleys, they seem to have held larger glacier systems in their catchments and may have more efficiently their residual SBSIS ice. 8.2 ka event 29 The largest cold excursion seen in the Holocene stable oxygen isotope record from the NGRIP ice core occurs at 8.2 ka BP marking the transition from Early to Mid Holocene (Young et al. 2012). The 8.2 ka BP event is believed to be a product of one of the catastrophic drainages from the collapsing Laurentide Ice Sheet which released a meltwater pulse from Hudson Bay into the Labrador Sea (Barber et al. 1999). These meltwater pulses or Heinrich Events are characterized by peaks in detrital carbonate seen in marine records through the North Atlantic (Heinrich 1988; Bond et al. 1993; Jennings et al. 2015). Studies have discussed the implication of meltwater pulses from the collapsing LIS influencing Svalbard glacier and climate during the Holocene (Hald & Korsun 2008; Hormes et al. 2013; van der Bilt et al. 2016). Despite the apparent climatic influence and subsequent glacial response seen on both sides of Baffin Bay to some of the peaks in detrital carbonate associated with the meltwater pulses (Young et al. 2012, 2013; Lesnek & Briner 2018), there is no terrestrial evidence yet identified that suggests similar glacier response in Svalbard. A high resolution marine sediment core from Van Mijenfjorden suggests reduction in δ18O of the benthic foraminifera C. reniforme as well as an increase in microfossil foraminifera implying a freshening and cooling of the fjord water around 8.2 ka BP. Although the catastrophic meltwater pulses may have influenced the oceanic conditions for a short time on Svalbard, the magnitude of ice margin retreat already experienced during the Early Holocene may have put the 8.2 ka BP ice fronts within the preceding Late Holocene glacier extent. Where glacial lake records have been used to reconstruct Holocene ice fluctuations for cirque glaciers, records suggest no peaks in mineragenic sedimentation suggesting glacial invigoration around 8.2 ka BP (Svendsen & Mangerud 1997; van der Bilt et al. 2016; de Wet et al. 2018; Røthe et al. 2018). We can cautiously interpret this lack of glacial evidence on Svalbard during the 8.2 event as a result of the magnitude of the Early Holocene ice-loss as well as the degree of Neoglacial-LIA re-advance. Whatever glacial response may have transpired as a result of this climatic event, it apparently occurred within the extent of the Late Holocene glacial maximum. Additionally, the high occurrence and wide diversity of thermosphilous molluscs during this transition from Early to Mid Holocene suggests the warm water species were seemingly unaffected by a mark cooling in fjord waters and Atlantic waters persisted widespread throughout Svalbard. Late Holocene (4.2 ka BP – present) The combination of decreasing summer insolation (Laskar et al. 2004) and the progressive cooling of regional waters around Svalbard draws the Late Holocene into the Neoglacial period (Fig. 10 & 11). 30 Marine- Trends in marine microfossil fauna assemblages indicate ocean cooling around Svalbard (Werner et al. 2013). Regional waters exhibit an increasing flux of IRD suggesting a growing glacial influence (Ślubowska-Woldengen et al. 2007; Rasmussen et al. 2014). Sedimentation rates in Svalbard fjords increases through the Late Holocene (Hald et al. 2004; Baeten et al. 2010; Kempf et al. 2013; Łącka et al. 2015; Streuff et al. 2017; Flink et al. 2017; Bartels et al. 2018). Reconstructions of sea ice cover derived from IP25 records spanning the Mid Holocene suggest a continued increase in ice-cover starting through the Late Holocene (Müller & Stein 2012; Bartels et al. 2018). In inner fjords, a decrease in IRD has been related to persistent sea ice suppressing iceberg rafting (Forwick and Vorren 2010). Fig. 10. Plot of 492 terrestrial SVALHOLA dates from the last 6 ka BP. Grey column delimits the end of the Mid Holocene. Specific radiocarbon ages from the database have been highlighted based on palaeo-glaciological and climatological associations. Dateable material reworked or over- 31 ridden by a glacier re-advance (blue), thermophilous marine molluscs (red) and ice cover expansion (green) are represented by enlarged colored circles with error bars. Histogram at base indicates TCN exposure age distribution within 500 year bins for moraine ridge boulders for three sites, L = Linnébreen, S = Scottbreen and T = Treskelen (Reusche et al. 2014; Philipps et al. 2017) Shaded bars at top represent a schematic view of Late Holocene glacial activity reconstructed with lake sediments of four glacial lakes from western Spitsbergen (Gjøavatnet, Hajeren, Kløsa and Linnévatnet). Blue boxes represent periods of glacial activity (dark blue suggest enhanced glacial activity), red boxes suggest no glacier was present in the catchment or glacial activity was greatly reduced; gray boxes indicate periods of reduced organic matter accumulation (Gjøavatnet; modified from de Wet et al. 2017). Narrow light blue boxes indicate periods of glacier growth interrupted by moraine stabilization (white boxes) reconstructed with lichenometry (Werner 1993). Terrestrial- Through the Late Holocene the ages of raised marine shorelines indicate a further decrease in relative uplift rates (Forman et al. 2004). Despite Mid Holocene transgression, it is believed that relative sea level was regressive throughout the Late Holocene. Currently, in several locations across northern and western Svalbard, uplift rates are surpassed by eustatic sea level and a transgression is being observed (Fjeldskaar et al. 2018). Although it is unclear when this modern transgression initiated, it is believed to be a relatively recent phenomenon, as coast line erosion proximal to pre-1900 cultural heritage and trapper huts have been some of the key indicators of active transgression (Andersson et al. 2000; Sessford et al. 2015). The period between 2.5 – 2.0 ka BP marks the Holocene minimum in driftwood arrival and matches the occurrence rate seen during the end of the Younger Dryas (Fig. 6). We assume the Holocene minimum in driftwood arrival relates to persistent or semi-permanent land-fast ice, minimizing the transport and catchment for driftwood accumulation (Funder et al. 2011). The increase in driftwood arrival exhibited through the last 2 ka of the Holocene record resembles the increase seen in the first half of the Early-Holocene (Fig. 6). Sea ice reconstructions derived from IP25 proxy records suggest variability in sea ice cover during the Late Holocene associated with sporadic warm sea surface temperatures (Müller et al. 2012; Sarnthein et al. 2003). Brine formation is believed to reflect sea ice cover above basins and a Late Holocene record from Storfjorden suggests episodic periods of intense production separated by periods of reduced brine formation (Rasmussen & Thomsen 2014). It is unclear if this trend-reversal in Holocene driftwood arrival is a result of decreasing sea ice cover around Svalbard (Fig. 6). The occurrence of Holocene driftwood on raised marine shorelines is influenced by preservation, and therefore may as well be influenced by Mid and Late Holocene shoreline transgressions. 32 The occurrence of thermophilous molluscs decreases through the latter half of the Holocene and tapers off entirely at the start of the Late Holocene (Fig. 10). In addition to the final two Mytilus edulis dating c. 3.7 ka BP, it appears the species may have returned to inner Isfjorden based on a single young sample dating 0.9 ka BP (Fig. 10; Samtleben 1985; Mangerud & Svendsen 2017). The lower occurrence of warm water molluscs through the Holocene suggests waters around Svalbard are cooling and the islands are transitioning into the Neoglacial in phase with declining summer insolation. The re-occurrence of the Mytilus at 0.9ka BP overlaps with the Medieval Warm Period and may suggest warm Atlantic waters were arriving to Svalbard. The mapped distribution of these warm water molluscs within fjords suggests the species thrive in the inner shallow branches of fjord-systems with minimal glacial influence (Dicksonfjorden and Billefjorden; Salvigsen et al. 1992; Salvigsen 2002; Blake et al. 2006; Mangerud & Svendsen 2017). An increasing number of glacier re-advances are constrained by radiocarbon dates of material re-sedimented in or overlain by glacial deposits throughout the Late Holocene (Fig. 10). These samples have been collected across Svalbard and become more frequent in the last 2 millennia (Fig. 6). The majority of the ages date to between 1.0 – 0.5 ka BP (Fig. 10). These sample dates constrain the maximum age of glacier re-advance during the Neoglacial or early LIA where they subsequently went on to reach their Late Holocene maximum extent (Baranowski & Karlén 1976; Dzierzek et al. 1990; Furrer 1991; Humlum et al. 2005). Several additional studies indicate Late Holocene glacier maxima occurred even earlier in the Neoglacial (Punning et al. 1976; Werner 1993; Sharin et al. 2014). Despite recent criticism, simplistic lichenometry studies suggest glaciers across Svalbard have experienced numerous phases of glacier advance, followed by moraine stabilization during the last 2.0 ka BP (Fig. 10; Werner 1993; Osborn et al. 2015). Exposure dating of moraine ridges has been conducted on several Late Holocene glacier forelands as a means of testing whether the LIA was the greatest glacial event of the Late Holocene (Fig. 10; Reusche et al. 2014; Philipps et al. 2017). The TCN studies have effectively supported earlier works, endorsing there have been numerous phases of glacial re-advance during the Late Holocene and concluding LIA is not the largest glacial event (at least in some locations; Werner 1993). Another technique suggesting an episodic phase of cooling during the Neoglacial is radiocarbon dating ice-entombed plants over a range of elevations to reconstruct snow-line- lowering or ice cover expansion throughout a region (Miller et al. 2017). Results from a study presenting over 40 radiocarbon ages from central Spitsbergen suggest that there were at least four phases of widespread ice cover expansion resulting in a general snow line lowering between 2.0 ka BP and 0.5 ka BP prior to the LIA-ice expansion (Miller et al. 2017). They further demonstrate that ice expansion occurred as early as 4.0 ka BP (Fig. 10; Miller et al. 2017). 33 Lacustrine- Late Holocene glacial lake records all suggest increasing glacial influence, characterized by increasing rates of sedimentation and transitions from organic-rich to mineragentic strata (Fig. 10). Other less maritime glacial lake records suggest episodic increases in glacial activity within catchments starting as early as c. 4 ka BP (Hajeren; van der Bilt et al. 2016) and prior to 3.0 ka BP (Kløsa and Linnévatnet; Røthe et al. 2015; Svendsen & Mangerud 1997). In Gjøavatnet (1 m a.s.l.), located in Northwestern Svalbard, the transition from organic-rich to glacial sedimentation does not occur until c. 1 ka BP. All glacial lake records show enhanced glacial influence during the last millennium (Fig. 10). Lacustrine records from non-glacial lakes (Hakluytvatnet and Skardsjøen) suggest decreasing productivity, declining plant species diversity as well as increased Late Holocene mineragenic input due to increased runoff (Alsos et al. 2015; Gjerde et al. 2018). Additionally, a 1.8 ka alkenone record from Kongressvatnet, located in central western Spitsbergen suggests the LIA was relatively mild in the region implying precipitation may be the key driver behind Late Holocene glacier expansion (D’Andrea et al. 2013). Neoglacial – Little Ice Age re-advances The classical perspective of the (Late) Holocene glacial maximum occurring during the culmination of the LIA is increasingly being challenged (Svendsen & Mangerud 1997; Snyder et al. 2000). Svalbard glaciers re-advanced throughout the Neoglacial period (Werner 1993; Reusche et al. 2014; Røthe et al. 2015; van der Bilt et al. 2016; Philipps et al. 2017) and during the LIA (Svendsen & Mangerud 1997; Snyder et al. 2000; Humlum et al. 2005; de Wet et al. 2018). Early observations of glaciers during the 1900s describe Svalbard glaciers near to their Late Holocene maximum extents. However this does not suggest anything about the timing of glacier re-advance, but rather that the glaciers were slowly starting to retreat. The majority of Late Holocene glacier re-advances (dated by overridden vegetation or mollusc shells) date to the early (between 1.0 - 0.5 ka BP), not the late LIA (Fig. 10). Several glacier re-advances constrained to the Neoglacial – LIA have been characterized as surges based on size, extent of glacial deposits and preservation of landforms (related to rapid ice advances) corresponding to associated ice-margins (Ottesen et al. 2008; Kristensen et al. 2009; Kempf et al. 2013; Farnsworth et al. 2016, 2017; Lovell & Boston 2017; Lyså et al. 2018; Flink et al. 2017). While the most extensive Late Holocene glacier deposits have been associated with surge-type behavior during, or at, the culmination of the LIA (Kristensen et al. 2009; Kempf et al. 2013; Flink et al. 2015; Lyså et al. 2018) an increasing number of studies have identified both 34 complete and fragmented moraine ridges outboard of the LIA maxima (Werner 1993; Sletten et al. 2001; Reusche et al. 2014; Philipps et al. 2017; Larsen et al. 2018). Little is known about patterns of Late Holocene precipitation across Svalbard. One alkenone-based (summer) temperature reconstruction from western Spitsbergen lake sediments suggests the Little Ice Age was “mild” and that precipitation played a larger contribution to regional glacier re-advances than previously acknowledged (D’Andrea et al. 2013). While it is important to highlight our limited knowledge of past precipitation, there is ample evidence suggesting that air and ocean temperatures were relatively cool and favored glaciers during the late Neoglacial and LIA (Divine et al. 2011; Bartels et al. 2017; Røthe et al. 2018; van der Bilt et al. 2018; Balascio et al. 2018; Luoto et al. 2018). Fig. 11 Timeline of Late Pleistocene and Holocene with divisions of the Holocene indicated (Cohen et al. 2013; updated). Curves of A) summer insolation at 80 °N (Wm-2; Laskar et al. 2004) B) the percentage of project glacier cover across Svalbard. Grey shade reflects uncertainty; low-end 35 estimates relate to evidence of the extent of small cirque and valley glaciers, while upper-end estimates correspond to high elevation ice caps. C) Reconstructed sea ice cover from Fram Strait derived from sea ice proxy IP25 (Müller et al. 2012; Müller & Stein 2014) and D) sea surface temperatures from western Spitsbergen derived from presence of warm water molluscs (Mangerud & Svendsen 2017) are presented. Sea ice conditions during the Neoglacial and LIA are characterized by increasing but variable ice extent around Svalbard (Fig. 11; Müller & Stein 2012; Müller et al. 2014 Bartels et al. 2017). Although increasing sea ice cover reconstructed through most of the Late Holocene is presumed to favor glacier growth (suppressing summer temperatures, minimizing the number of positive degree days and decreasing frontal ablation) it is unclear how much this restricts precipitation. A summary timeline of reconstructed conditons from Svalbard indicates the relationship of estimated glacier cover over Svalbard compared reconstructed sea ice conditions and sea surface temperature as well as summer insolation (Fig. 11; Laskar et al. 2004). Glaciers were most likely at their maximum Holocene extent during the onset of the Holocene. The Mid Holocene is characterized by a low percentage ice cover. Glaciers re-advance during the Early and Late Holocene. The marine Holocene thermal optimum occurred in the Early Holocene while the terrestrial thermal optium likely occurred during the Mid Holocene (Mangerud & Svendsen 2017). Glaciers located in small, low elevation valleys and cirques likely degalciated completely during the Mid Holocene (de Wet et al. 2018). Marine terminating outlet glaciers seem to remain throughout the Mid Holocene in some locations. However, these outlet glaciers likely corresponded to high elevation catchments. Glaciers re-advanced through the Late Holocene and predominantly reach Neoglacial maximums some time in the last millennium. Widespread glacier retreat has been exhibited across Svalbard through the last century. SUMMARY (1) Is there evidence of Younger Dryas cooling and glacier re-advances on Svalbard? No evidence of a synchronous glacial event is observed during the Younger Dryas on Svalbard. There is mounting evidence suggesting glacier re-advances during the end of the Younger Dryas into the Early Holocene. These deposits are not well dated, but re-advances appear to span the transition from Late Pleistocene to Early Holocene. This phase of glacier activity may relate to glacio-dynamic behavior resulting from a transition into a period with less extensive sea ice cover leading to increased precipitation as well as warming oceans and atmosphere. These factors all have the potential to influence glacier hydrological systems and thermal, which could result in a 36 dynamic, unsustainable advance or surge-type behavior (Dunse et al. 2015; Østby et al. 2017; Farnsworth et al. 2018; Sevestre et al. 2015, 2018). (2) Did meltwater from the collapsing Laurentide Ice Sheet influence climate on Svalbard (8.2 ka BP)? It is possible, but data remains inconclusive. In select locations, marine stratigraphic studies suggest cooling and freshening of waters around 8.2 ka BP. We are unable to know if these shifts in ocean conditions were great enough to impact the Svalbard glaciers systems. The absence of a Svalbard glacier event relating to the 8.2 ka BP meltwater pulse may reflect both the severity of Early Holocene retreat and the magnitude of Late Holocene re-advances during the Neoglacial and Little Ice Age. As of yet, there has been no 8.2 ka BP glacial event identified on Svalbard. Less is known about the influence of earlier melt water pulses from the LIS. (3) When was the Holocene glacial minimum and what was the ice cover during this period? We do not (yet) know when glaciers on Svalbard reached a Holocene minimum, nor how reduced ice cover became. While lake sediment studies targeting cirque glaciers suggest that catchments were ice-free through the Mid Holocene, lake studies have yet to target larger glacier systems. Marine IRD studies suggest that tidewater glaciers in some fjords retreated back to unknown terrestrial positions, while other fjords probably hosted a calving ice front throughout the entire Holocene. The Holocene glacial minimum probably occurred in the Mid Holocene sometime between 8.0 - 6.5 ka BP. The glacial minimum certainly occurred after the marine thermal optimum (10.0 – 9.2 ka BP), potentially lagging the peak in ocean temperatures and summer insolation on a millennium time scale. (4) When was the most extensive ice cover during the Holocene period and was the Little Ice Age the climax of the Neoglacial? Ice cover was probably most extensive at the onset of the Holocene based on the presumed amount of residual SBSIS-ice. Glacier re-advances have been identified, both in the Early and Late Holocene, but it is unknown which phase of glacier re-advance was the greatest. The LIA is generally believed to be the biggest Holocene glacial event on Svalbard. However, increasing numbers of moraines outboard of the LIA extent of both Neoglacial and Early Holocene age suggest ice margins extended further at different periods during the Holocene. As we do not yet know how reduced glaciers were prior to the advance, it is difficult to reconcile which event was of 37 the greatest magnitude. Furthermore the lack of knowledge regarding the rate of advance limits our understanding of how rapidly Svalbard glaciers have responded to shifts in Holocene climate. (5) Which sedimentary archives provide the best detail of Holocene glaciers and climate? This depends on which questions one is targeting. But studies should have a holistic approach to systems and processes; taking into account a mixture of methods, archives and disciplines. (6) To what extend has ice dynamics and surge-type behavior influenced Holocene glacier fluctuations on Svalbard? Ice dynamics play an important role in glacier fluctuations today. The extent is unknown, but ice dynamics are believed to have also played a role in ice marginal fluctuations during the Early and the Late Holocene. • Throughout the Holocene, Svalbard glaciers have responded to a varying combination of climatic, environmental and dynamic driving factors which influence both the extent and behavior of ice margins. • Glaciers during the Late Pleistocene and Early Holocene were dynamic, exhibited re- advances and extended well beyond Late Holocene glacier maxima in many locations across Svalbard. • The marine Holocene thermal optimum on Svalbard marked by the early arrival of warm water species to the coasts of Svalbard during the onset of the Holocene, pre-dates the peak in northern hemisphere summer insolation and the terrestrial thermal optimum. • Glaciers covered a small, but unknown percentage of Svalbard, in some locations terminated at sea level, during the Holocene glacial minimum. • Evidence of episodic Neoglacial glacier re-advances is being identified more commonly across Svalbard, suggesting an irregular phase of glacier favorable conditions during the Late Holocene. • The Little Ice Age maximum does not reflect the glacial maximum extent across Svalbard; since the deglaciation, the onset of the Holocene or even during the Late Holocene, in most locations. 38 REFERENCES for SVALHOLA Database Aagaard-Sørensen et al. 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Geografiska Annaler Series A 72, 301–317. 187 Appendix Chapter I -Fig S1 Topographic maps of Lateglacial Early Holocene ice margins -Fig S2 Aerial images of Lateglacial Early Holocene ice margins -Table S1 Glacier Length Data -Table S2 Lateglacial-early Holocene studies -References Chapter II -Table S1 Radiocarbon dates from Sveagruva area -Quaternary geological & geomorphological map Svea, Svalbard (link) Chapter III -Not applicable Chapter IV -Table S1 Documented surge-type glaciers -Table S2 Glaciers with crevasse squeeze ridges Chapter V -Table S1 Radiocarbon ages from the SVALHOLA database (link) -Table S2 Luminesence ages from the SVALHOLA database (link) -Table S3 Cosmogenic exposure ages from the SVALHOLA database (link) -Table S4 References from SVALHOLA database Figure S1. Topographic maps of Lateglacial-early Holocene (LGEH) glacier deposits labeled A-F corresponding to glaciers; (A) Tåbreen, (B) Albertbreen, (C) Lyckholmdalen, (D) Flowerbreen, (E) Heftybreen and (F) Richterbreen respectively. Light blue shades represent inferred modern glacier extent, Late Holocene (LH) max. extent and Lateglacial early Holocene min. extent. Remnant glacier deposits indicated with dotted lines while dark shade of blue represents high relative sea level. Inset map exhibits location of respective sites on Spitsbergen. Topographic maps modified from TopoSvalbard (http://toposvalbard.npolar.no/). Figure S2. Enlarged high resolution aerial imagery of six selected sites presented in Fig. 4, where glacier deposits (marked by white dotted lines) are remotely constrained in age to Lateglacial-early Holocene age. Imagery freely accessible on TopoSvalbard provided by the Norwegian Polar Institute. Table S1. Glacier Length Data ID Region Location Glacier Length (km)* Snout to NG (km) NG –LGEH (km) Snout –LGEH (km) Glacier type - Lomfjorden De Geerbukta Gullfaksebreen 19.64 1.2 4.37 5.57 outlet (Åsgardfonna) A Wijdefjorden Mosseldalen Tåbreen 7.63 0.29 3.81 4.1 outlet (Åsgardfonna) B Liefdefjorden Sørdalsbukta Albertbreen 2.61 0.47 2.30 2.77 cirque glacier C Dicksonfjorden Lyckholmdalen Lyckholmbreen* 1.90 0.61 8.10 8.71 outlet (Friggkåpa) D Sassenfjorden Vindodden Flowerbreen 2.76 1.01 4.09 5.1 valley glacier E Grønfjorden Sandefjorden Heftybreen 1.34 1.02 2.65 3.67 cirque glacier F Van Keulenfjorden Richter/Ulladalen Richterbreen 5.04 1.32 4.15 5.47 valley glacier Note: Distance measurements made in TopoSvalbard (2017). *Unofficial glacier names based valley or site name. Abbreviations are as follows; NG = Neoglacial; LGEH = Lateglacial-early Holcoene TABLE S2. Lateglacial-early Holocene studies Ref. ID Reference Location Glacier / Site Focus Sample 14C age BP Cal. a BP Median Description 1 Eitel et al. (2002) Andreeland Gråhukdalen Morphology LGEH advance 2 Eitel et al. (2002) Andreeland Vogtdalen Stratigraphy / Map Shell frag. 10096±77 11022 Max. constraining age 3 Eitel et al. (2002) Andreeland Vatnedalen Morphology LGEH advance 4 Eitel et al. (2002) Andreeland Cirrusdalen Morphology LGEH advance 5 Eitel et al. (2002) Andreeland Torrelldalen Stratigraphy / Map Shell frag. 10340±120 11310 Max. constraining age 6 Eitel et al. (2002) Andreeland Jakobsenbukta Stratigraphy / Map Shell frag. 10580±60 11722 Max. constraining age 7 Brückner et al. (2002) Andreeland Verdalspynten Stratigraphy Shell 9311±57 10074 Min. constraining age (Hd 20997) 8 Eitel et al. (2002) Andreeland Wigdehlpynten Morphology LGEH advance 9 Salvigsen & Österholm (1982) Reinsdyrflya Sørdalsflya Morphology Lateglacial ice margin suggested 10 van der Bilt et al. (2015) Mitrahalvøya Hajeren Lake sediments Terrestrial Macro Age-depth model - Early Holocene glacial maxima 11 Røthe et al. (2015) Mitrahalvøya Kløsa / Karlbreen Lake sediments Terrestrial Macro Age-depth model - Pre 6.7 ka BP 12 Henriksen et al. (2014) Kongsfjorden Kongsfjordhallet Stratigraphy / Map Moraine Boulder 9200 - Cosmogenic nuclide 12 Henriksen et al. (2014) Kongsfjorden Kongsfjordhallet Stratigraphy / Map Moraine Boulder 12100 - Cosmogenic nuclide 13 Andersson et al. (1999) Prins Karls Forland McVitiepynten Stratigraphy / Map Shell 12590±70 14002 Max. constraining age (T-2096) 14 Andersson et al. (2000) Prins Karls Forland Poolepynten Stratigraphy / Map Whalebone 11650±180 13054 Max. constraining age (1-13795) 14 Farnsworth et al. (2017) St. Jonsfjorden Gunnarbreen Stratigraphy / Map N.labradorica (Foram) 11345±102 12750 Max. constraining age (HH12.956) 15 Farnsworth et al. (2017) St. Jonsfjorden Gunnarbreen Stratigraphy / Map M. modiolous 9320±60 10081 Min. constraining age (LuS 10795) 16 Salvigsen et al. 1990) Isfjorden Ymerbukta Stratigraphy / Map M. truncata 9500±100 10290 Max. constraining age (T-6286) 17 Lønne 2005) Bolterdalen Scott Turnerbreen Stratigraphy / Map M. truncata 9775±125 10595 Max. constraining age (T-13883) 17 Lønne 2005) Bolterdalen Scott Turnerbreen Stratigraphy / Map M. truncata 9625±95 10408 Min. constraining age (T-13884) 18 Landvik & Salvigsen 1985) Reindalen Gangdalen Stratigraphy - - - Post deglaciation advance 19 Lønne 2005) (Salvigsen) Grønfjorden Aldegondabreen Stratigraphy Shell frag. 9980±120 10871 Max. constraining age (T-6290) 20 Larson et al. (in press) Van Mijenfjorden Gustravdalen Stratigraphy / Map 10340±110 11392 Max. constraining age (Tln-146 3,6) 20 Larson et al. (in press) Van Mijenfjorden Gustravdalen Shoreline Shell 9510±90 10367 Min. constraining age (Tln-145 3,6) 21 Mangerud et al. (1992) Van Mijenfjorden Bromelldalen Stratigraphy / Map Shell 10940±200 12319 Max. constraining age (T-5368) 21 Mangerud et al. (1992) Van Mijenfjorden Bromelldalen Stratigraphy / Map Shell 10570±150 11659 Min. constraining age (T-5369) 22 Landvik et al. (1992) Bellsund Skilvika Stratigraphy Shell Pre LGM tributary advance 23 Ronnert & Landvik (1993) Edgeøya Albrechtbreen Stratigraphy / Map Whalebone 8635±125 9211 Max. constraining age (T-9911) - van der Bilt et al. (2018) Mitrahalvøya / Amsterdamøya Hajeren / Hakluytvatnet Lake sed. / Alkenones Terrestrial macros Age-depth model - Early warmth 12-10.5 ka BP - Balascio et al. (2018) Amsterdamøya Hakluytvatnet Lake sed. / Leaf waxes Terrestrial macros Age-depth model - Early warmth 12.8-9.5 ka BP - Mangerud & Svendsen (2018) Spitsbergen Isfjorden Thermophiles Molluscs Review with 21 new ages - Early warm fjords 11-9 ka BP - Hald et al. (2004) Spitsbergen West Spitsbergen Marine Sed. / Proxy Forams Age-depth model - Early warm regional waters 11.2-8.8 ka BP - Müller & Stein (2014) Spitsbergen West Spitsbergen Marine Sed. / Proxy IP 25 Age-depth model - Sea-ice reduction ~12 ka BP Note: All dates corrected for marine reservoir age with a ΔR = 70 ± 30 (Mangerud & Svendsen 2017) and calibrated with Marine13 on Calib Rev 7.0.4 References: Andersen, K. K., Azuma, N., Barnola, J. M., & Bigler, M. 2004: High-resolution record of Northern Hemisphere climate extending into the last interglacial period. Nature 431(7005), 147. Andersson, T., Forman, S. L., Ingólfsson, Ó. & Manley, W. F. 1999: Late Quaternary environmental history of central Prins Karls Forland, western Svalbard. Boreas. 28, 292-307. Andersson, T., Forman, S. L., Ingólfsson, Ó. & Manley, W. F. 2000: Stratigraphic and Morphologic Constraints on the Weichselian Glacial History of Northern Prins Karls Forland, Western Svalbard. Geografiska Annaler A 82, 455-470. Balascio, N. L., D'Andrea, W. J., Gjerde, M. & Bakke, J. 2018: Hydroclimate variability of High Arctic Svalbard during the Holocene inferred from hydrogen isotopes of leaf waxes. Quaternary Science Reviews doi.org/10.1016/j.quascirev.2016.11.036. van der Bilt, W. G. M., Bakke, J., Vasskog, K., D'Andrea, W. J., Bradley, R. S. & Ólafsdóttir, S. 2015: Reconstruction of glacier variability from lake sediments reveals dynamic Holocene climate in Svalbard. Quaternary Science Reviews 126, 201-218. van der Bilt, W. G. M., D'Andrea, W. J., Bakke, J., Balascio, N. L., Werner, J. P., Gjerde, M. & Bradley, R. S. 2018: Alkenone-based reconstructions reveal four-phase Holocene temperature evolution for High Arctic Svalbard. 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Landvik, J. Y., Bolstad, M., Lycke, A. K., Mangerud, J. & Sejrup, H. P. 1992: Weichselian stratigraphy and palaeoenvironments at Bellsund, western Svalbard. Boreas 21, 335-358. Lønne, I. 2005: Faint traces of high Arctic glaciations: an early Holocene ice-front fluctuation in Bolterdalen, Svalbard. Boreas 34, 308-323. Mangerud, J., Bolstad, M., Elgersma, A., Helliksen, D., Landvik, J. Y., Lønne, I., Lycke, A. K., Salvigsen, O., Sandahl, T. & Svendsen, J. I. 1992: The last glacial maximum on Spitsbergen, Svalbard. Quaternary Research 38, 1-31. Mangerud, J. & Svendsen, J. I. 2018: The Holocene Thermal Optimum around Svalbard, Arctic North Atlantic; molluscs show early and exceptional warmth. The Holocene XX, 1-19. Müller, J. & Stein, R. 2014: High-resolution record of late glacial and deglacial sea ice changes in Fram Strait corroborates ice–ocean interactions during abrupt climate shifts. Earth and Planetary Science Letters 403, 446-455. Reimer, P. J., Bard, E., Bayliss, A., Beck, J. W., Blackwell, P. G., Bronk Ramsey, C., Buck, C. E., Cheng, H., Edwards, R. L., Friedrich, M., Grootes, P. M., Guilderson, T. P., Haflidason, H., Hajdas, I., Hatté, C., Heaton, T. J., Hoffmann, D. L., Hogg, A. G., Hughen, K. A., Kaiser, K. F., Kromer, B., Manning, S. W., Niu, M., Reimer, R. W., Richards, D. A., Scott, E. M., Southon, J. R., Staff, R. A., Turney, C. S. M. & van der Plicht, J. 2013: IntCal13 and Marine13 Radiocarbon Age Calibration Curves 0–50,000 Years cal BP. Radiocarbon 55, 1869-1887 Ronnert, L. & Landvik, J. Y. 1993: Holocene glacial advances and moraine formation at Albrechtbreen, Edgeøya, Svalbard. Polar Research 12, 57-63. Røthe, T. O., Bakke, J., Vasskog, K., Gjerde, M., D'Andrea, W. J. & Bradley, R. S. 2015: Arctic Holocene glacier fluctuations reconstructed from lake sediments at Mitrahalvøya, Spitsbergen. Quaternary Science Reviews 109, 111-125. Salvigsen, O., Elgersma, A., Hjort, C., Lagerlund, E., Liestøl, O. & Svensson, N.-O. 1990: Glacial history and shoreline displacement on Erdmannflya and Bohemanflya, Spitsbergen, Svalbard. Polar Research 8, 261-273. Salvigsen, O. & Österholm, H. 1982: Radiocarbon dated raised beaches and glacial history of the northern coast of Spitsbergen, Svalbard. Polar Research 1, 97-115. Documented surge-type glaciers Supplementary Table 1 Glacier Name Island Region Lat. / Long. CSR Photo year Ref. Nord Buchananisen Prince Karls Foreland NWS 78.6863 N, 11.0047 E Yes 2008 J. et al. 2000 Aavatsmarkbreen Spitsbergen NWS 78.7000 N, 12.1467 E Yes 2009 B. et al. 2009 Abrahamsenbreen Spitsbergen NWS 79.1050 N, 14.4097 E Yes 2011 Liestøl 1993 Ankerbreen Spitsbergen NWS 78.5499 N, 12.6494 E Yes 2009 Hamilton 1992 Anna Sofiebreen Spitsbergen NWS 78.4899 N, 13.1787 E No 2009 Hamilton 1992 Arbobreen Spitsbergen NWS 78.9744 N, 15.6381 E Yes 2011 Hamilton 1992 Austgotabreen Spitsbergen NWS 78.4375 N, 13.0506 E Yes 2009 Hamilton 1992 Austre Broggerbreen Spitsbergen NWS 78.8889 N, 11.8537 E Yes 2010 Hamilton 1992 Battyebreen Spitsbergen NWS 78.9731 N, 15.0005 E Yes 2011 Hamilton 1992 Biksebreen Spitsbergen NWS 78.8769 N, 14.9838 E Yes 2011 Hamilton 1992 Blindernbreen Spitsbergen NWS 78.9867 N, 15.2820 E Yes 2011 Hamilton 1992 Blomstrandbreen Spitsbergen NWS 79.0652 N, 12.3765 E Yes 2009 Liestøl 1993 Blåshaugbreen Spitsbergen NWS 79.1967 N, 11.4738 E No 2009 Hamilton 1992 Borebreen Spitsbergen NWS 78.4422 N, 13.8830 E Yes 2009 Hamilton 1992 Botnfjellbreen Spitsbergen NWS 78.8429 N, 12.4199 E Yes 2009 Hamilton 1992 Bukkebreen Spitsbergen NWS 78.5449 N, 13.4647 E Yes 2009 Hamilton 1992 Charlesbreen Spitsbergen NWS 78.5068 N, 13.3628 E Yes 2009 Hamilton 1992 Chauveauxbreen Spitsbergen NWS 79.6424 N, 11.9868 E No 2009 S. et al. 2009 Cissybreen Spitsbergen NWS 78.7479 N, 11.9397 E Yes 2009 Hamilton 1992 Comfortlessbreen Spitsbergen NWS 78.7635 N, 12.1304 E Yes 2009 S. et al. 2009 Delbreen Spitsbergen NWS 78.8738 N, 15.7890 E Yes 2011 Hamilton 1992 Eidembreen Spitsbergen NWS 78.4065 N, 13.1986 E Yes 2009 Hamilton 1992 Elisebreen Spitsbergen NWS 78.6467 N, 12.2550 E Yes 2009 S. et al. 2009 Elnabreen Spitsbergen NWS 79.1596 N, 14.1118 E Yes 2011 Liestøl 1993 Erikkabreen Spitsbergen NWS 78.7330 N, 12.0671 E Yes 2009 Hamilton 1992 Fjortende Julibreen Spitsbergen NWS 79.1224 N, 12.1576 E Yes 2009 M. et al., 2012 Fyrisbreen Spitsbergen NWS 78.9384 N, 15.5715 E Yes 2011 Liestøl 1993 Gaffelbreen Spitsbergen NWS 78.5671 N, 12.8878 E Yes 2009 Hamilton 1992 Gonvillebreen Spitsbergen NWS 78.7754 N, 16.0328 E No 2011 Hamilton 1992 Gyldenbreen Spitsbergen NWS 78.8580 N, 16.0815 E Yes 2011 Hamilton 1992 Haakenbreen Spitsbergen NWS 78.7452 N, 11.9759 E Yes 2009 Hamilton 1992 Hansdalsbreen Spitsbergen NWS 78.6899 N, 15.0196 E No 2011 Hamilton 1992 Høegdalsbreen Spitsbergen NWS 78.9516 N, 15.7087 E Yes 2011 Hamilton 1992 Holmstrombreen Spitsbergen NWS 78.8318 N, 14.0566 E Yes 2009 Liestøl 1993 Hørbyebreen Spitsbergen NWS 78.7623 N, 16.2847 E Yes 2011 Hamilton 1992 Hydrografbreen Spitsbergen NWS 78.4544 N, 12.7489 E Yes 2009 Hamilton 1992 Kappfjellbreen Spitsbergen NWS 79.0840 N, 11.9651 E Yes 2009 Hamilton 1992 Kastellbreen Spitsbergen NWS 78.9730 N, 15.9871 E No 2011 Hamilton 1992 Kinamurbreen Spitsbergen NWS 78.8152 N, 15.9694 E Yes 2011 Hamilton 1992 Klampebreen Spitsbergen NWS 78.6016 N, 13.4796 E Yes 2009 Hamilton 1992 Kongsvegen Spitsbergen NWS 78.8154 N, 12.7764 E Yes 2009 Liestøl 1993 Konowbreen Spitsbergen NWS 78.6025 N, 12.9480 E Yes 2009 Hamilton 1992 Konsulbreen Spitsbergen NWS 79.0961 N, 12.3668 E No 2009 Hamilton 1992 Kronebreen Spitsbergen NWS 78.8747 N, 12.7883 E Yes 2009 Hamilton 1992 Lisbetbreen Spitsbergen NWS 79.0679 N, 15.0314 E Yes 2011 Liestøl 1993 Manchesterbreen Spitsbergen NWS 78.8148 N, 16.0747 E Yes 2011 Hamilton 1992 Mannbreen Spitsbergen NWS 78.9506 N, 15.8503 E Yes 2011 Hamilton 1992 Mediumbreen Spitsbergen NWS 78.5379 N, 14.1445 E No 2009 Hamilton 1992 Meyerbreen Spitsbergen NWS 78.6969 N, 15.0265 E Yes 2011 Hamilton 1992 Midtre Lovénbreen Spitsbergen NWS 78.8782 N, 12.0415 E Yes 2009 Liestøl 1993 Monacobreen Spitsbergen NWS 79.3979 N, 12.6080 E No 2009 B. et al. 2009 Muninbreen Spitsbergen NWS 78.7245 N, 16.0895 E Yes 2011 Hamilton 1992 Nonsenbreen Spitsbergen NWS 78.3747 N, 13.8966 E Yes 2009 Liestøl 1993 Nordenfjeldskebreen Spitsbergen NWS 78.7905 N, 12.2445 E Yes 2009 Hamilton 1992 Nordre Vasskilbreen Spitsbergen NWS 78.9142 N, 15.6626 E Yes 2011 Hamilton 1992 Olssonbreen Spitsbergen NWS 79.0600 N, 11.9472 E No 2009 Hamilton 1992 Orsabreen Spitsbergen NWS 78.9688 N, 14.2764 E Yes 2011 J. et al. 2000 Osbornebreen Spitsbergen NWS 78.6550 N, 13.1276 E No 2009 D. et al. 1991 Paulbreen Spitsbergen NWS 78.5603 N, 13.4415 E Yes 2009 Hamilton 1992 Roysbreen Spitsbergen NWS 78.9049 N, 13.4415 E Yes 2009 Hamilton 1992 Samebreen Spitsbergen NWS 78.6337 N, 14.1172 E No 2009 Hamilton 1992 Sefstrombreen Spitsbergen NWS 78.7468 N, 13.9464 E Yes 2009 Liestøl 1993 Skansdalsbreen Spitsbergen NWS 78.5588 N, 15.7619 E Yes 2011 Liestøl 1993 Smalfjellbreen Spitsbergen NWS 78.7751 N, 12.2156 E No 2009 Hamilton 1992 Studentbreen Spitsbergen NWS 78.9696 N, 15.4662 E Yes 2011 Hamilton 1992 Størmerbreen Spitsbergen NWS 79.0281 N, 14.5840 E Yes 2009 S. et al. 2009 Supanbreen Spitsbergen NWS 79.3288 N, 11.9827 E Yes 2009 Hamilton 1992 Sveabreen Spitsbergen NWS 78.6937 N, 13.6096 E Yes 2009 Hamilton 1992 Svelgfjellet Nord* Spitsbergen NWS 79.2921 N, 14.5535 E Yes 2010 Liestøl 1993 Svenbreen Spitsbergen NWS 78.7259 N, 16.3060 E Yes 2011 Hamilton 1992 Tindebreen Spitsbergen NWS 79.6995 N, 11.9562 E No 2009 Hamilton 1992 Uversbreen Spitsbergen NWS 78.8107 N, 12.3480 E Yes 2009 Hamilton 1992 Vegardbreen Spitsbergen NWS 78.5319 N, 13.4192 E Yes 2009 Hamilton 1992 Vestgotabreen Spitsbergen NWS 78.4611 N, 12.9015 E Yes 2009 Hamilton 1992 Vestre Lovenbreen Spitsbergen NWS 78.9028 N, 11.9216 E Yes 2010 Hamilton 1992 Vintervegen Spitsbergen NWS 78.5846 N, 13.4866 E Yes 2009 S. et al. 2009 Wahlenbergbreen Spitsbergen NWS 78.5825 N, 13.8586 E Yes 2009 Liestøl 1993 Allfarvegen Spitsbergen NES 79.0535 N, 19.7140 E Yes 2011 L & H 1991 Backlundbreen Spitsbergen NES 78.6458 N, 19.9593 E Yes 2011 L & H 1991 Ebbabreen Spitsbergen NES 78.7287 N, 16.8639 E Yes 2011 Hamilton 1992 Eskolabreen Spitsbergen NES 78.9149 N, 16.6148 E Yes 2011 Hamilton 1992 Ganskijbreen Spitsbergen NES 78.7210 N, 20.6709 E Yes 2011 L & H 1991 Hannbreen Spitsbergen NES 78.7684 N, 21.2900 E Yes 2011 L & H 1991 Hayesbreen Spitsbergen NES 78.3756 N, 18.5826 E No 2011 Liestøl 1993 Hochstetterbreen Spitsbergen NES 78.8810 N, 20.4419 E Yes 2011 Liestøl 1993 Kantbreen Spitsbergen NES 79.3508 N, 17.8594 E Yes 2010 J. et al. 2000 Kosterbreen Spitsbergen NES 79.3175 N, 18.7222 E Yes 2010 L & H 1991 Longstaffbreen Spitsbergen NES 79.7762 N, 16.0648 E No 2010 Liestøl 1993 Mittag Lefflerbreen Spitsbergen NES 78.8768 N, 16.4652 E Yes 2011 Hamilton 1992 Negribreen Spitsbergen NES 78.5929 N, 18.7927 E Yes 2011 L & H 1991 Nordenskïoldbreen Spitsbergen NES 78.6580 N, 17.0094 E Yes 2009 J. et al. 2000 Odinjøkulen Spitsbergen NES 79.5746 N, 18.3008 E No 2010 L & H 1991 Oslobreen Spitsbergen NES 79.0905 N, 18.7590 E No 2011 Liestøl 1993 Pedasenkobreen Spitsbergen NES 78.7071 N, 21.0479 E Yes 2011 L & H 1991 Petermannbreen Spitsbergen NES 78.4993 N, 18.7832 E Yes 2011 J. et al. 2000 Sonklarbreen Spitsbergen NES 78.7240 N, 20.3340 E Yes 2011 L & H 1991 Tunabreen Spitsbergen NES 78.4739 N, 17.4255 E Yes 2009 Liestøl 1993 Vestre Odinjokulen* Spitsbergen NES 79.5792 N, 18.1914 E No 2010 Liestøl 1993 Von Postbreen Spitsbergen NES 78.4209 N, 17.5214 E Yes 2009 Liestøl 1993 Andrinebreen Spitsbergen CS 77.9573 N, 17.0216 E No 2011 S. et al. 2009 Ankerbreen Spitsbergen CS 77.9566 N, 16.3181 E Yes 2011 Hamilton 1992 Ayerbreen Spitsbergen CS 78.0888 N, 16.0147 E Yes 2009 Hamilton 1992 Bergmesterbreen Spitsbergen CS 78.0797 N, 16.9182 E Yes 2009 Hamilton 1992 Blackbreen Spitsbergen CS 78.2859 N, 16.1383 E No 2009 Hamilton 1992 Blekumbreen Spitsbergen CS 78.2478 N, 16.0051 E No 2009 Hamilton 1992 Brandtbreen Spitsbergen CS 78.2868 N, 15.9483 E No 2009 Hamilton 1992 Bødalen Spitsbergen CS 78.0953 N, 16.4452 E No 2011 Hamilton 1992 Drønbreen Spitsbergen CS 78.1319 N, 16.8223 E Yes 2009 Hamilton 1992 Duboisbreen Spitsbergen CS 78.2601 N, 16.7753 E No 2009 Hamilton 1992 Elfenbeinbreen Spitsbergen CS 78.1643 N, 18.1371 E No No Hamilton 1992 Fangenbreen Spitsbergen CS 78.2887 N, 16.2126 E Yes 2009 Hamilton 1992 Fridtjovbreen Spitsbergen CS 77.7953 N, 14.5101 E No 2008 Liestøl 1993 Glitrebreen Spitsbergen CS 78.0295 N, 17.1445 E No 2009 Hamilton 1992 Gløttfjellbreen Spitsbergen CS 78.1391 N, 16.4566 E No 2009 Hamilton 1992 Helsingborgbreen Spitsbergen CS 77.9597 N, 16.8318 E Yes 2009 Hamilton 1992 Höganäsbreen Spitsbergen CS 77.9398 N, 16.7633 E No 2009 Hamilton 1992 Innerbreen Spitsbergen CS 78.1057 N, 17.4057 E No 2011 Hamilton 1992 Klauvbreen Spitsbergen CS 78.2997 N, 18.1949 E No 2012 Hamilton 1992 Klunsbreen Spitsbergen CS 78.1563 N, 18.2662 E No 2012 Hamilton 1992 Kokbreen Spitsbergen CS 78.0801 N, 16.4633 E Yes 2009 Hamilton 1992 Livbreen Spitsbergen CS 77.9375 N, 16.3441 E Yes 2009 Hamilton 1992 Lunckebreen Spitsbergen CS 78.0315 N, 16.7935 E Yes 2009 Hamilton 1992 Margitbreen Spitsbergen CS 78.0074 N, 17.7975 E No 2011 Hamilton 1992 Marmorbreen Spitsbergen CS 78.1777 N, 17.8231 E No 2012 Hamilton 1992 Marthabreen Spitsbergen CS 78.0193 N, 16.6904 E Yes 2009 Hamilton 1992 Medalsbreen Spitsbergen CS 78.0249 N, 15.5412 E No 2011 Hamilton 1992 Møysalbreen Spitsbergen CS 78.1341 N, 16.4935 E No 2009 Hamilton 1992 Målarbreen Spitsbergen CS 78.2709 N, 16.0848 E Yes 2009 Hamilton 1992 Pålsjöbreen Spitsbergen CS 77.9992 N, 17.0231 E Yes 2009 Hamilton 1992 Scott Turnerbreen Spitsbergen CS 78.1076 N, 15.9598 E Yes 2009 Hamilton 1992 Skoltbreen Spitsbergen CS 78.1421 N, 16.7374 E No 2009 Hamilton 1992 Skruisbreen Spitsbergen CS 78.1524 N, 17.8133 E Yes 2012 Hamilton 1992 Skutbreen Spitsbergen CS 78.0355 N, 17.5192 E Yes 2011 Hamilton 1992 Sveigbreen Spitsbergen CS 78.1093 N, 17.8000 E Yes 2011 Hamilton 1992 Svellnosbreen Spitsbergen CS 78.0802 N, 16.4094 E No 2009 Hamilton 1992 Tillbergfonna Spitsbergen CS 78.0717 N, 15.7960 E Yes 2011 Hamilton 1992 Tinkarpbreen Spitsbergen CS 77.9876 N, 17.1855 E No 2009 S. et al. 2009 Ulvebreen Spitsbergen CS 78.2096 N, 18.6497 E Yes 2012 Hamilton 1992 Usherbreen Spitsbergen CS 78.2727 N, 18.6724 E Yes 2012 Hamilton 1992 Vegbreen Spitsbergen CS 78.0250 N, 16.9524 E Yes 2009 Hamilton 1992 Vendombreen Spitsbergen CS 78.1429 N, 17.2921 E No 2009 Hamilton 1992 Amundsenisen Spitsbergen SS 77.2842 N, 15.5406 E No 2010 S. et al. 2009 Anna Margrethebreen Spitsbergen SS 77.3415 N, 17.3157 E Yes 2011 L & H 1991 Arnesenbreen Spitsbergen SS 77.7986 N, 18.0908 E Yes 2012 L & H 1991 Aurkollfonna Spitsbergen SS 77.3983 N, 15.9905 E No 2011 Hamilton 1992 Austjøkulen Spitsbergen SS 76.8761 N, 16.6199 E Yes 2010 S. et al. 2009 Bakaninbreen Spitsbergen SS 77.7590 N, 17.4863 E Yes 2011 D. et al. 1991 Barettbreen Spitsbergen SS 77.1845 N, 16.6386 E No 2011 Hamilton 1992 Beinbreen Spitsbergen SS 77.1176 N, 16.9637 E No 2011 Hamilton 1992 Bellingbreen Spitsbergen SS 77.3100 N, 17.3039 E Yes 2011 Hamilton 1992 Bendfjellbreen Spitsbergen SS 77.2137 N, 16.6902 E No 2011 Hamilton 1992 Bereznikovbreen Spitsbergen SS 77.7927 N, 18.2338 E Yes 2012 L & H 1991 Billesholmbreen Spitsbergen SS 77.8982 N, 17.1603 E Yes 2009 Hamilton 1992 Bjuvbreen Spitsbergen SS 77.9046 N, 17.2194 E Yes 2009 Hamilton 1992 Blümckebreane Spitsbergen SS 77.5287 N, 15.0317 E No 2011 Hamilton 1992 Bosarpbreen Spitsbergen SS 77.8814 N, 17.0558 E Yes 2009 Hamilton 1992 Bratthengbreen Spitsbergen SS 77.7560 N, 18.3776 E No 2012 Hamilton 1992 Brodtkorbfjellet Spitsbergen SS 77.8393 N, 17.1965 E Yes 2011 S. et al. 2009 Bungebreen Spitsbergen SS 76.8153 N, 16.0941 E Yes 2010 S. et al. 2009 Buttbreen Spitsbergen SS 77.6846 N, 16.4830 E Yes 2011 Hamilton 1992 Charpentierbreen Spitsbergen SS 77.6163 N, 15.5409 E Yes 2011 Liestøl 1993 Crollbreen Spitsbergen SS 77.2070 N, 17.2646 E Yes 2011 L & H 1991 Davisbreen Spitsbergen SS 77.2463 N, 17.2500 E Yes 2011 Liestøl 1993 Dobrowolskibreen Spitsbergen SS 77.3499 N, 16.6929 E No 2011 S. et al. 2009 Doktorbreen Spitsbergen SS 77.5276 N, 16.6780 E Yes 2011 Hamilton 1992 Dvergfonna Spitsbergen SS 77.4603 N, 15.4428 E No 2011 Hamilton 1992 Edvardbreen Spitsbergen SS 77.8885 N, 17.6170 E Yes 2011 S. et al. 2009 Filantropbreen Spitsbergen SS 77.7115 N, 16.7107 E Yes 2011 Hamilton 1992 Finsterwalderbreen Spitsbergen SS 77.4595 N, 15.2446 E Yes 2011 Hamilton 1992 Firmbreen Spitsbergen SS 77.8703 N, 17.2276 E No 2009 S. et al. 2009 Flatbreen Spitsbergen SS 77.1553 N, 16.8051 E Yes 2011 Hamilton 1992 Fredfonna Spitsbergen SS 76.9126 N, 16.6501 E No 2010 S. et al. 2009 Gribnefjellbreen Spitsbergen SS 77.7558 N, 17.6153 E No 2011 Hamilton 1992 Hambergbreen Spitsbergen SS 77.0534 N, 17.0115 E Yes 2011 L & H 1991 Hessbreen Spitsbergen SS 77.5079 N, 15.1125 E Yes 2011 Liestøl 1993 Hettebreen Spitsbergen SS 77.7472 N, 16.7945 E Yes 2011 Hamilton 1992 Hornbreen Spitsbergen SS 77.0640 N, 16.6945 E Yes 2011 J. et al. 2000 Hyllingebreen Spitsbergen SS 77.9090 N, 17.2928 E Yes 2009 Liestøl 1993 Høgstebereen Spitsbergen SS 77.3334 N, 15.1421 E Yes 2011 S. et al. 2009 Indrebøbreen Spitsbergen SS 77.6123 N, 17.9269 E Yes 2012 S. et al. 2009 Ingerbreen Spitsbergen SS 77.7481 N, 18.2846 E Yes 2012 S. et al. 2009 Inglefieldbreen Spitsbergen SS 77.8263 N, 17.9488 E Yes 2012 L & H 1991 Isbroddbreen Spitsbergen SS 77.2087 N, 16.9720 E No 2011 Hamilton 1992 Karibreen Spitsbergen SS 77.6089 N, 17.7828 E Yes 2012 Hamilton 1992 Keipbreen Spitsbergen SS 77.7741 N, 17.6222 E No 2011 Hamilton 1992 Søre Kjølhø* Spitsbergen SS 77.6375 N, 17.2950 E No 2012 S. et al. 2009 Kjølhøbreen Spitsbergen SS 77.6647 N, 17.3053 E No 2012 S. et al. 2009 Kleivbreen Spitsbergen SS 77.9115 N, 17.3469 E Yes 2009 S. et al. 2009 Klubbebreen Spitsbergen SS 77.7225 N, 17.0614 E Yes 2011 S. et al. 2009 Knoppbreen Spitsbergen SS 77.6781 N, 17.2340 E No 2012 S. et al. 2009 Kroppbreen Spitsbergen SS 77.8967 N, 17.4149 E Yes 2011 S. et al. 2009 Liten Kroppebreen* Spitsbergen SS 77.9154 N, 17.5779 E No 2011 S. et al. 2009 Kvalbreen Spitsbergen SS 77.5822 N, 18.0621 E Yes 2012 S. et al. 2009 Kvastbreen Spitsbergen SS 77.4456 N, 17.2824 E Yes 2011 Hamilton 1992 Körberbreen Spitsbergen SS 76.9463 N, 16.0814 E Yes 2010 Liestøl 1993 Langleikbreen Spitsbergen SS 77.2170 N, 16.5182 E No 2011 Hamilton 1992 Langryggbreen Spitsbergen SS 77.3859 N, 15.9164 E No 2011 Hamilton 1992 Lognbreen Spitsbergen SS 77.7793 N, 17.3864 E No 2011 Hamilton 1992 Luntebreen Spitsbergen SS 77.7240 N, 16.8725 E No 2011 Hamilton 1992 Markhambreen Spitsbergen SS 77.1211 N, 17.2016 E Yes 2011 S. et al. 2009 Martinbreen Spitsbergen SS 77.6446 N, 15.6694 E Yes 2011 Liestøl 1993 Mendelejevbreen Spitsbergen SS 76.9276 N, 16.5523 E Yes 2010 S. et al. 2009 Mühlbacherbreen Spitsbergen SS 77.1461 N, 15.9167 E No 2011 S. et al. 2009 Noglebreen Spitsbergen SS 77.5568 N, 17.2035 E No 2012 S. et al. 2009 Notaschabreen Spitsbergen SS 77.6898 N, 17.5438 E No 2011 S. et al. 2009 Nothorstbreen Spitsbergen SS 77.2833 N, 16.7208 E Yes 2011 S. et al. 2009 Nordsysselbreen Spitsbergen SS 77.8493 N, 17.7867 E No 2011 Hamilton 1992 Nornebreen Spitsbergen SS 77.2288 N, 15.7106 E No 2011 S. et al. 2009 Novbreen Spitsbergen SS 77.2071 N, 16.7680 E No 2011 Hamilton 1992 Oksbreen Spitsbergen SS 77.1311 N, 16.9866 E No 2011 Hamilton 1992 Paierlbreen Spitsbergen SS 77.1229 N, 15.7230 E No 2011 B. et al. 2009 Paulabreen Spitsbergen SS 77.7362 N, 17.3125 E Yes 2011 S. et al. 2009 Peisbreen Spitsbergen SS 77.7381 N, 17.1280 E Yes 2011 Hamilton 1992 Penckbreen Spitsbergen SS 77.4741 N, 15.6391 E Yes 2011 Hamilton 1992 Persejbreen Spitsbergen SS 77.4575 N, 17.3803 E Yes 2012 D. et al. 2003 Plogbreen Spitsbergen SS 76.7866 N, 16.1875 E Yes 2010 Hamilton 1992 Polakkbreen Spitsbergen SS 77.2729 N, 16.1693 E No 2011 S. et al. 2009 Profilbreen Spitsbergen SS 77.2769 N, 15.1777 E Yes 2011 S. et al. 2009 RagNo-Mariebreen Spitsbergen SS 77.7942 N, 17.3597 E Yes 2011 S. et al. 2009 Recherchebreen Spitsbergen SS 77.4370 N, 14.8208 E No 2011 Liestøl 1993 Richardsbreen Spitsbergen SS 77.7314 N, 18.0373 E Yes 2012 B. et al. 2009 Richterbreen Spitsbergen SS 77.6429 N, 15.5267 E Yes 2011 Hamilton 1992 Ryggbreen Spitsbergen SS 77.1666 N, 17.0547 E No 2011 Hamilton 1992 Samarinbreen Spitsbergen SS 76.8858 N, 16.3490 E Yes 2010 Hamilton 1992 Scheelebreen Spitsbergen SS 77.6929 N, 16.9702 E Yes 2011 S. et al. 2009 Scottbreen Spitsbergen SS 77.5439 N, 14.3668 E Yes 2011 Liestøl 1993 Sergievskifjellet Spitsbergen SS 77.5357 N, 17.4598 E Yes 2012 S. et al. 2009 Siegerbreen Spitsbergen SS 77.4513 N, 15.8743 E Yes 2011 Liestøl 1993 Skilfonna Spitsbergen SS 76.9307 N, 16.9131 E Yes 2010 S. et al. 2009 Skimebreen Spitsbergen SS 77.2834 N, 17.2410 E Yes 2011 B. et al. 2009 Skobreen Spitsbergen SS 77.7062 N, 17.2081 E No 2012 S. et al. 2009 Snokuvbreen Spitsbergen SS 77.7021 N, 16.3589 E Yes 2011 S. et al. 2009 Spaelbreen Spitsbergen SS 77.4186 N, 17.2275 E Yes 2011 Hamilton 1992 St. Olgafjellet Spitsbergen SS 77.5830 N, 17.7880 E No 2012 S. et al. 2009 Stabbfonna Spitsbergen SS 77.1466 N, 16.9560 E No 2011 Hamilton 1992 Staupbreen Spitsbergen SS 77.0813 N, 17.2518 E Yes 2011 Liestøl 1993 Steenstrupbreen Spitsbergen SS 77.6053 N, 16.2919 E Yes 2011 Hamilton 1992 Stolbreen Spitsbergen SS 77.1360 N, 17.1078 E Yes 2011 Hamilton 1992 Storbreen Spitsbergen SS 77.1299 N, 16.3370 E Yes 2010 Hamilton 1992 Strongbreen Spitsbergen SS 77.5815 N, 17.5069 E Yes 2012 L & H 1991 Sulsbreen Spitsbergen SS 77.8155 N, 17.7378 E No 2011 Hamilton 1992 Svalbreen Spitsbergen SS 77.6608 N, 16.5462 E Yes 2011 Hamilton 1992 Sykorabreen Spitsbergen SS 77.0179 N, 17.0045 E Yes 2011 L & H 1991 Søkkbreen Spitsbergen SS 77.6204 N, 17.3235 E No 2012 S. et al. 2009 Thomsonbreen Spitsbergen SS 77.6380 N, 18.2098 E Yes 2012 L & H 1991 Tjørndalsbreen Spitsbergen SS 77.5362 N, 14.2363 E No 2011 Hamilton 1992 Vallåkrabreen (Bringen) Spitsbergen SS 77.8600 N, 17.1127 E Yes 2009 S. et al. 2009 Vasilievbreen Spitsbergen SS 76.7894 N, 16.7383 E Yes 2010 L & H 1991 Vasilievbreen (Rundtuva) Spitsbergen SS 76.7023 N, 16.8632 E No 2010 L & H 1991 Venetzbreen Spitsbergen SS 77.6016 N, 15.5816 E Yes 2011 Hamilton 1992 Vindeggbreen Spitsbergen SS 77.4932 N, 17.2880 E Yes 2012 S. et al. 2009 Zawadskibreen Spitsbergen SS 77.3390 N, 15.8828 E Yes 2011 S. et al. 2009 Bodleybreen Nord Austlandet NA 79.8370 N, 21.4901 E No 2011 D. et al. 1999 Bråsvellbreen (Af B1) Nord Austlandet NA 79.3680 N, 23.4389 E No 2011 D. et al. 1999 Austfonna (B3) Nord Austlandet NA 79.7769 N, 24.6528 E No 2011 Liestøl 1993 Clasebreen Nord Austlandet NA 79.5775 N, 20.3272 E No 2011 L & H 1991 Franklinbreen nordre Nord Austlandet NA 80.1050 N, 19.6123 E No 2010 D. et al. 1999 Franklinbreen Søre Nord Austlandet NA 80.0629 N, 19.4651 E No 2010 Liestøl 1993 Palanderbreen Nord Austlandet NA 79.5320 N, 21.0316 E No 2011 L & H 1991 Palenderisen Nord Austlandet NA 79.5665 N, 21.8560 E No 2011 L & H 1991 Rijpbreen Nord Austlandet NA 80.1013 N, 21.5677 E No 2010 Liestøl 1993 Etonbreen Nord Austlandet NA 79.7298 N, 21.9163 E Yes 2011 D. et al. 1999 Holtenbreen Nord Austlandet NA 79.5642 N, 20.4385 E No 2011 L & H 1991 Duckwitzbreen Barentsøya BEØ 78.3416 N, 20.8540 E Yes 2010 Liestøl 1993 Freemanbreen Barentsøya BEØ 78.2911 N, 21.7744 E Yes 2010 Liestøl 1993 Hübnerbreen Barentsøya BEØ 78.3718 N, 21.9848 E Yes 2010 Liestøl 1993 Reymondbreen Barentsøya BEØ 78.3943 N, 22.0054 E Yes 2010 Liestøl 1993 Austre Kvitisen Edgeøya BEØ 78.0531 N, 21.6311 E Yes 2010 Liestøl 1993 Deltabreen Edgeøya BEØ 77.5590 N, 23.0057 E Yes 2010 L & H 1991 Kong Johans Bre Edgeøya BEØ 77.6268 N, 23.8996 E Yes 2010 Liestøl 1993 Kuhrbreen Edgeøya BEØ 77.5567 N, 21.7792 E Yes 2010 L & H 1991 Marsjøbreen Edgeøya BEØ 77.9367 N, 22.4144 E Yes 2010 Liestøl 1993 Pettersenbreen Edgeøya BEØ 77.4809 N, 23.4976 E Yes 2010 Liestøl 1993 Skrentbreen Edgeøya BEØ 77.5363 N, 21.5107 E Yes 2010 Liestøl 1993 Stonebreen Edgeøya BEØ 77.8110 N, 23.9041 E Yes 2010 L & H 1991 Suraustre Bergfonna Edgeøya BEØ 78.1171 N, 22.0333 E No 2010 Liestøl 1993 Vestre Kvitkåpa Edgeøya BEØ 77.4043 N, 22.8036 E Yes 2010 Liestøl 1993 B. et al. 2009 Blaszczyk et al. 2009 D. et al. 1991 Dowdeswell et al. 1991 D. et al. 1999 Dowdeswell et al. 1999 D. et al. 2003 Dowdeswell et al. 2003 Hamilton 1992 Hamiliton 1992 J. et al. 2000 Jiskoot et al. 2000 L & H 1991 Lefauconnier and Hagen 1991 Liestøl 1993 Liestøl 1993 S. et al. 2009 Sund et al. 2009 Glaciers with CSR Supplementary Table 2 Glacier Name Island Region Lat. / Long. CSR Photo year Ref. Alfredbreen Prince Karls Foreland NWS 78.5753 N, 11.1083 E Yes 2008 This study Archibald Geikiebreen Prince Karls Foreland NWS 78.4532 N, 11.5596 E Yes 2008 This study Doddsbreen Prince Karls Foreland NWS 78.5188 N, 11.3129 E Yes 2008 This study Fallbreen Prince Karls Foreland NWS 78.7011 N, 10.9903 E Yes 2008 This study James Geikiebreen Prince Karls Foreland NWS 78.4955 N, 11.4219 E Yes 2008 This study Magdabreen Prince Karls Foreland NWS 78.5654 N, 11.1672 E Yes 2008 This study Midtre Buchananisen Prince Karls Foreland NWS 78.6641 N, 11.0177 E Yes 2008 This study Millerbreen Prince Karls Foreland NWS 78.7085 N, 10.8261 E Yes 2008 This study Murraybreen Prince Karls Foreland NWS 78.7292 N, 11.0023 E Yes 2008 This study Parnassbreen Prince Karls Foreland NWS 78.6275 N, 10.9316 E Yes 2008 This study Søre Buchananisen Prince Karls Foreland NWS 78.6444 N, 11.0670 E Yes 2008 This study Agnorbreen Spitsbergen NWS 78.6571 N, 12.1823 E Yes 2009 This study Albertbreen Spitsbergen NWS 79.6578 N, 12.7306 E Yes 2011 This study Andreasbreen Spitsbergen NWS 78.6123 N, 12.2633 E Yes 2009 This study Andreebreen Spitsbergen NWS 79.6692 N, 12.4784 E Yes 2011 This study Angelbreen Spitsbergen NWS 79.0704 N, 15.3702 E Yes 2011 This study Anne-Mariebreen Spitsbergen NWS 79.3924 N, 14.9266 E Yes 2010 This study Arthurbreen Spitsbergen NWS 78.7722 N, 11.9098 E Yes 2009 This study Aulbreen Spitsbergen NWS 78.3213 N, 13.2483 E Yes 2009 This study Austre Holmesletbreen Spitsbergen NWS 78.4896 N, 12.9385 E Yes 2009 This study Austre Lovenbreane Spitsbergen NWS 78.8746 N, 12.1463 E Yes 2009 This study Balliolbreen Spitsbergen NWS 78.8067 N, 16.3140 E Yes 2011 This study Bardebreen Spitsbergen NWS 78.6739 N, 14.1829 E Yes 2009 This study Barmfjellbreen Spitsbergen NWS 78.9068 N, 14.7812 E Yes 2011 This study Belshornbreen Spitsbergen NWS 79.1683 N, 15.0329 E Yes 2011 This study Bertilbreen Spitsbergen NWS 78.6912 N, 16.2658 E Yes 2011 This study Binnebreen Spitsbergen NWS 79.2771 N, 15.2119 E Yes 2010 This study Bitihornbreen Spitsbergen NWS 78.8703 N, 14.6593 E Yes 2011 This study Botnefjellet Spitsbergen NWS 78.8586 N, 12.4051 E Yes 2009 This study Brurskankbreen Spitsbergen NWS 78.9641 N, 14.8508 E Yes 2011 This study Bukkebreen Spitsbergen NWS 79.1258 N, 15.1951 E Yes 2011 This study Bullbreen Spitsbergen NWS 78.4867 N, 12.6591 E Yes 2009 This study Børrebreen Spitsbergen NWS 79.5118 N, 13.0568 E Yes 2011 This study Caiusbreen Spitsbergen NWS 78.7645 N, 16.0552 E Yes 2011 This study Cambridgebreen Spitsbergen NWS 78.8001 N, 16.3845 E Yes 2011 This study Conwaybreen Spitsbergen NWS 78.9929 N, 12.6170 E Yes 2009 This study Dahlbreen Spitsbergen NWS 78.6021 N, 12.5327 E Yes 2009 This study D'Arodesbreen Spitsbergen NWS 79.1537 N, 12.0970 E Yes 2009 This study Eddabreen Spitsbergen NWS 79.2717 N, 14.8895 E Yes 2010 This study Edithbreen Spitsbergen NWS 78.8569 N, 12.0977 E Yes 2009 This study Eivindbreen Spitsbergen NWS 78.6310 N, 12.2802 E Yes 2009 This study Erikbreen Spitsbergen NWS 79.6187 N, 12.4779 E Yes 2011 This study Ernstbreen Spitsbergen NWS 79.5849 N, 14.4412 E Yes 2010 This study Esmarkbreen Spitsbergen NWS 78.3205 N, 13.8122 E Yes 2009 This study Evabreen Spitsbergen NWS 79.6870 N, 12.6347 E Yes 2011 This study Feiringbreen Spitsbergen NWS 79.0199 N, 12.4807 E Yes 2009 This study Ferdinandbreen Spitsbergen NWS 78.7104 N, 16.3408 E Yes 2011 This study Flakbreen Spitsbergen NWS 79.1740 N, 11.9764 E Yes 2009 This study Førstebreen Spitsbergen NWS 79.2706 N, 11.1494 E Yes 2009 This study Garmbreen Spitsbergen NWS 79.7505 N, 14.5524 E Yes 2010 This study Gavlhaugbreen Spitsbergen NWS 78.8760 N, 15.5155 E Yes 2011 This study Geabreen Spitsbergen NWS 78.3295 N, 14.0666 E Yes 2009 This study Georgbreen Spitsbergen NWS 79.5662 N, 14.4422 E Yes 2010 This study Ginnungagapbreen Spitsbergen NWS 79.2136 N, 14.6479 E Yes 2010 This study Gislebreen Spitsbergen NWS 78.4758 N, 12.5535 E Yes 2009 This study Glopbreen Spitsbergen NWS 79.5284 N, 13.0110 E Yes 2011 This study Gufsbreen Spitsbergen NWS 78.8651 N, 14.8311 E Yes 2011 This study Gullmarbreen Spitsbergen NWS 79.6722 N, 11.3604 E Yes 2011 This study Gunnarbreen Spitsbergen NWS 78.4908 N, 13.1080 E Yes 2009 This study Gygrebreen Spitsbergen NWS 78.8054 N, 14.9144 E Yes 2011 This study Harrietbreen Spitsbergen NWS 78.2667 N, 13.5973 E Yes 2009 This study Hillbreen Spitsbergen NWS 79.6188 N, 11.2314 E Yes 2011 This study Hodsbreen Spitsbergen NWS 78.9883 N, 15.5275 E Yes 2011 This study Hoelbreen Spitsbergen NWS 78.7552 N, 16.2600 E Yes 2011 This study Holmesletbreane Spitsbergen NWS 78.4891 N, 12.9020 E Yes 2009 This study Hønnbreen Spitsbergen NWS 79.1612 N, 15.1796 E Yes 2011 This study Høvbreen Spitsbergen NWS 79.4369 N, 14.4308 E Yes 2010 This study Isrypebreen Spitsbergen NWS 78.7033 N, 14.0743 E Yes 2009 This study Johanbreen Spitsbergen NWS 79.1642 N, 13.9657 E Yes 2011 This study Kaalaasbreen Spitsbergen NWS 79.0198 N, 15.6983 E Yes 2011 This study Karlsbreen Spitsbergen NWS 79.3468 N, 13.4088 E Yes 2011 This study Keisarbreen Spitsbergen NWS 79.5532 N, 12.9111 E Yes 2011 This study Kjerulfbreen Spitsbergen NWS 78.2873 N, 13.6556 E Yes 2009 This study Kongsbreen Spitsbergen NWS 78.9255 N, 12.7588 E Yes 2009 This study Kyrkjebreen Spitsbergen NWS 78.8747 N, 14.0060 E Yes 2009 This study Landbreen Spitsbergen NWS 79.6815 N, 14.5829 E Yes 2010 This study Lappbreen Spitsbergen NWS 78.6412 N, 14.2767 E Yes 2009 This study Lexfjellbreen Spitsbergen NWS 78.2617 N, 13.4925 E Yes 2009 This study Lilliehöökbreen Spitsbergen NWS 79.3501 N, 11.7191 E Yes 2009 This study Løvliebreen Spitsbergen NWS 78.4846 N, 13.0154 E Yes 2009 This study Maribreen Spitsbergen NWS 78.9498 N, 15.4010 E Yes 2011 This study Miethebreen Spitsbergen NWS 79.5628 N, 11.2986 E Yes 2011 This study Morabreen Spitsbergen NWS 78.8694 N, 14.2196 E Yes 2009 This study Målarbreen Spitsbergen NWS 78.8520 N, 14.5902 E Yes 2011 This study Ottobreen Spitsbergen NWS 79.6089 N, 14.4820 E Yes 2010 This study Pedersenbreen Spitsbergen NWS 78.8682 N, 12.2868 E Yes 2009 This study Protektorbreen Spitsbergen NWS 78.2379 N, 13.6791 E Yes 2009 This study Purpurbreen Spitsbergen NWS 78.9585 N, 15.9234 E Yes 2011 This study Qvarnstrombreen Spitsbergen NWS 78.7485 N, 14.3778 E Yes 2009 This study Ringertzbreen Spitsbergen NWS 79.2047 N, 13.5174 E Yes 2011 This study Robertsonbreen Spitsbergen NWS 78.7902 N, 15.9749 E Yes 2011 This study Sagtindbreen Spitsbergen NWS 79.2802 N, 12.2090 E Yes 2009 This study Schjelderupbreen Spitsbergen NWS 79.3955 N, 13.3686 E Yes 2011 This study Sellströmbreen Spitsbergen NWS 79.6960 N, 11.3319 E Yes 2011 This study Serlabreen Spitsbergen NWS 79.7307 N, 12.6631 E Yes 2011 This study Skaugumbreen Spitsbergen NWS 79.5556 N, 14.7806 E Yes 2010 This study Skreifjellbreen Spitsbergen NWS 79.0171 N, 12.2921 E Yes 2009 This study Slørbreen Spitsbergen NWS 79.2835 N, 11.4513 E Yes 2009 This study Smalgangenbreen Spitsbergen NWS 78.5617 N, 13.0116 E Yes 2009 This study Smithbreen Spitsbergen NWS 79.7382 N, 11.8618 E Yes 2011 This study Southhamptonbreen Spitsbergen NWS 78.8082 N, 16.1549 E Yes 2011 This study Stabeisbreen Spitsbergen NWS 79.0065 N, 15.0167 E Yes 2011 This study Stallobreen Spitsbergen NWS 78.3746 N, 13.2644 E Yes 2009 This study Svardalsbreen Spitsbergen NWS 79.6774 N, 14.8455 E Yes 2010 This study Svelgfjellbreen Søre Spitsbergen NWS 79.2855 N, 14.4713 E Yes 2010 This study Sølvbreen Spitsbergen NWS 78.8482 N, 14.8794 E Yes 2011 This study Søre Vasskilbreen Spitsbergen NWS 78.8922 N, 15.6646 E Yes 2011 This study Tassbreen Spitsbergen NWS 78.7830 N, 11.8956 E Yes 2009 This study Torgnybreen Spitsbergen NWS 78.2380 N, 13.5723 E Yes 2009 This study Uggbreen Spitsbergen NWS 79.2333 N, 15.1339 E Yes 2010 This study Ujamnbreen Spitsbergen NWS 78.5781 N, 12.6174 E Yes 2009 This study Universitetsbreen Spitsbergen NWS 78.9846 N, 15.3607 E Yes 2011 This study Venbreen Spitsbergen NWS 78.9362 N, 14.7178 E Yes 2011 This study Venernbreen Spitsbergen NWS 78.3403 N, 13.2113 E Yes 2009 This study Veobreen Spitsbergen NWS 79.2257 N, 11.4782 E Yes 2009 This study Vestre Brøggerbreane Spitsbergen NWS 78.9131 N, 11.7426 E Yes 2009 This study Vetternbreen Spitsbergen NWS 78.2924 N, 13.3811 E Yes 2009 This study Vogtbreen Spitsbergen NWS 79.6590 N, 14.6250 E Yes 2010 This study Vonbreen Spitsbergen NWS 79.2006 N, 13.7021 E Yes 2011 This study Waldemarbreen Spitsbergen NWS 78.6749 N, 12.0469 E Yes 2009 This study Yggbreen Spitsbergen NWS 79.1096 N, 15.2932 E Yes 2011 This study * NWS-1 Spitsbergen NWS 79.5936 N, 14.7008 E Yes 2010 This study * NWS-2 Spitsbergen NWS 79.5234 N, 14.6465 E Yes 2010 This study * NWS-3 Spitsbergen NWS 79.5028 N, 14.3789 E Yes 2010 This study * NWS-4 Spitsbergen NWS 79.4849 N, 14.6117 E Yes 2010 This study * NWS-5 Spitsbergen NWS 79.3715 N, 14.5618 E Yes 2010 This study * NWS-6 Spitsbergen NWS 79.3317 N, 14.7224 E Yes 2010 This study * NWS-7 Spitsbergen NWS 79.3044 N, 14.8876 E Yes 2010 This study * NWS-8 Spitsbergen NWS 79.2418 N, 14.5704 E Yes 2011 This study * NWS-9 Spitsbergen NWS 79.1827 N, 12.2420 E Yes 2009 This study * NWS-10 Spitsbergen NWS 79.1417 N, 14.7231 E Yes 2011 This study * NWS-11 Spitsbergen NWS 79.1245 N, 14.6855 E Yes 2011 This study * NWS-12 Spitsbergen NWS 79.1056 N, 14.6841 E Yes 2011 This study * NWS-13 Spitsbergen NWS 79.0271 N, 14.8898 E Yes 2011 This study * NWS-14 Spitsbergen NWS 79.0009 N, 15.6862 E Yes 2011 This study * NWS-15 Spitsbergen NWS 78.8666 N, 15.9476 E Yes 2011 This study * NWS-16 Spitsbergen NWS 78.8311 N, 14.8404 E Yes 2011 This study * NWS-17 Spitsbergen NWS 78.8005 N, 11.8869 E Yes 2009 This study * NWS-18 Spitsbergen NWS 78.7973 N, 11.8371 E Yes 2009 This study * NWS-19 Spitsbergen NWS 78.7752 N, 16.5404 E Yes 2011 This study * NWS-20 Spitsbergen NWS 78.7137 N, 16.1528 E Yes 2011 This study * NWS-21 Spitsbergen NWS 78.7117 N, 14.9811 E Yes 2011 This study * NWS-22 Spitsbergen NWS 78.7070 N, 15.9753 E Yes 2011 This study * NWS-23 Spitsbergen NWS 78.7034 N, 15.7074 E Yes 2011 This study * NWS-24 Spitsbergen NWS 78.7025 N, 15.8999 E Yes 2011 This study * NWS-25 Spitsbergen NWS 78.6914 N, 16.0344 E Yes 2011 This study * NWS-26 Spitsbergen NWS 78.6320 N, 15.8399 E Yes 2011 This study * NWS-27 Spitsbergen NWS 78.6169 N, 14.2329 E Yes 2009 This study * NWS-28 Spitsbergen NWS 78.6135 N, 16.0110 E Yes 2011 This study * NWS-29 Spitsbergen NWS 78.6034 N, 12.3576 E Yes 2009 This study * NWS-30 Spitsbergen NWS 78.5577 N, 12.4685 E Yes 2009 This study * NWS-31 Spitsbergen NWS 78.5374 N, 15.8909 E Yes 2011 This study * NWS-32 Spitsbergen NWS 78.5243 N, 15.7647 E Yes 2011 This study * NWS-33 Spitsbergen NWS 78.5240 N, 14.2120 E Yes 2009 This study * NWS-34 Spitsbergen NWS 78.5150 N, 14.2284 E Yes 2009 This study * NWS-35 Spitsbergen NWS 78.4665 N, 14.1794 E Yes 2009 This study * NWS-36 Spitsbergen NWS 78.4660 N, 13.9817 E Yes 2009 This study * NWS-37 Spitsbergen NWS 78.4625 N, 14.1963 E Yes 2009 This study * NWS-38 Spitsbergen NWS 78.4308 N, 12.6925 E Yes 2009 This study * NWS-39 Spitsbergen NWS 78.3853 N, 14.0697 E Yes 2009 This study * NWS-40 Spitsbergen NWS 78.3593 N, 13.1976 E Yes 2009 This study * NWS-41 Spitsbergen NWS 78.3414 N, 13.8878 E Yes 2009 This study * NWS-42 Spitsbergen NWS 78.2963 N, 13.8004 E Yes 2009 This study * NWS-43 Spitsbergen NWS 78.2908 N, 13.8159 E Yes 2009 This study Beckerfjelletbreen Spitsbergen NES 78.9795 N, 20.2595 E Yes 2011 This study Berglibreen Spitsbergen NES 79.6698 N, 16.0693 E Yes 2010 This study Bogebreen Spitsbergen NES 78.4607 N, 17.5689 E Yes 2009 This study Brucebreen Spitsbergen NES 78.4671 N, 17.2855 E Yes 2009 This study Buldrebreen Spitsbergen NES 79.8088 N, 17.6740 E Yes 2010 This study Burn Murdochbreen Spitsbergen NES 78.4590 N, 17.0617 E Yes 2009 This study Chydeniusbreen Spitsbergen NES 79.2379 N, 18.2541 E Yes 2010 This study Dunerbreen Spitsbergen NES 79.8115 N, 16.8123 E Yes 2010 This study Dvergbreen Spitsbergen NES 79.4602 N, 17.7084 E Yes 2010 This study Fairweatherbreen Spitsbergen NES 78.5796 N, 16.8609 E Yes 2009 This study Florabreen Spitsbergen NES 78.5975 N, 17.2908 E Yes 2009 This study Formidablebreen Spitsbergen NES 78.8771 N, 16.5887 E Yes 2011 This study Frøyabreen Spitsbergen NES 79.5469 N, 18.1483 E Yes 2010 This study Glintbreen Spitsbergen NES 79.3756 N, 17.9801 E Yes 2010 This study Glærbreen Spitsbergen NES 79.6187 N, 17.5781 E Yes 2010 This study Gullfaksebreen Spitsbergen NES 79.5629 N, 17.4294 E Yes 2010 This study Heuglinbreen Spitsbergen NES 78.3794 N, 18.8252 E Yes 2012 This study Hinlopenbreen Spitsbergen NES 79.0721 N, 18.9968 E Yes 2011 This study Hodbreen Spitsbergen NES 79.3616 N, 18.7956 E Yes 2010 This study Hønerbreen Spitsbergen NES 79.1259 N, 18.7711 E Yes 2011 This study Johansenbreen Spitsbergen NES 78.5244 N, 18.7531 E Yes 2011 This study Keiserkampen Spitsbergen NES 79.0486 N, 20.6421 E Yes 2011 This study Koristkabreen Spitsbergen NES 78.8191 N, 21.2386 E Yes 2011 This study Königsbreen Spitsbergen NES 78.3439 N, 18.6524 E Yes 2012 This study Laubefjellebreen Spitsbergen NES 78.8501 N, 21.0229 E Yes 2011 This study Loderbreen Spitsbergen NES 79.2031 N, 18.8506 E Yes 2011 This study McWhaebreen Spitsbergen NES 78.7792 N, 16.6587 E Yes 2011 This study Methuenbreen Spitsbergen NES 78.5249 N, 17.1335 E Yes 2009 This study Nordbreen Spitsbergen NES 79.6378 N, 15.9158 E Yes 2010 This study Ordonnansbreen Spitsbergen NES 78.6390 N, 18.9873 E Yes 2011 This study Planckbreen Spitsbergen NES 79.1834 N, 16.4026 E Yes 2010 This study Polarisbreen Spitsbergen NES 79.2260 N, 18.5248 E Yes 2011 This study Ragnarbreen Spitsbergen NES 78.7540 N, 16.7125 E Yes 2011 This study Reinsbukkbreen Spitsbergen NES 79.2025 N, 16.6432 E Yes 2010 This study Reliktbreen Spitsbergen NES 79.4365 N, 18.7668 E Yes 2010 This study Rimfaksebreen Spitsbergen NES 79.6007 N, 17.5133 E Yes 2010 This study Ringhornbreen Spitsbergen NES 79.3324 N, 16.4563 E Yes 2010 This study Roonbreen Spitsbergen NES 79.0132 N, 19.7522 E Yes 2011 This study Royal Societybreen Spitsbergen NES 79.3115 N, 16.4657 E Yes 2010 This study Sanderbreen Spitsbergen NES 79.0894 N, 16.5005 E Yes 2011 This study Sistebreen Spitsbergen NES 79.3703 N, 17.3138 E Yes 2010 This study Skinfaksbebreen Spitsbergen NES 79.5038 N, 17.3460 E Yes 2010 This study Stubendorffbreen Spitsbergen NES 78.9623 N, 16.6063 E Yes 2011 This study Sven Ludvigbreen Spitsbergen NES 79.3868 N, 18.7860 E Yes 2010 This study Sørbreen Spitsbergen NES 79.4944 N, 15.9824 E Yes 2010 This study Tommelbreen Spitsbergen NES 79.4965 N, 18.6292 E Yes 2010 This study Tryggvebreen Spitsbergen NES 79.1081 N, 16.5216 E Yes 2011 This study Veitebreen Spitsbergen NES 79.1546 N, 18.8300 E Yes 2011 This study Veteranbreen Spitsbergen NES 79.3335 N, 17.3172 E Yes 2010 This study Wichefjelletbreen Spitsbergen NES 78.4673 N, 18.9005 E Yes 2011 This study * NES-1 Spitsbergen NES 79.6632 N, 17.8133 E Yes 2010 This study * NES-2 Spitsbergen NES 79.2160 N, 18.8140 E Yes 2011 This study * NES-3 Spitsbergen NES 78.6239 N, 19.3801 E Yes 2011 This study * NES-4 Spitsbergen NES 78.4485 N, 18.9460 E Yes 2011 This study * NES-5 Spitsbergen NES 78.4241 N, 18.9752 E Yes 2011 This study Aldegondabreen Spitsbergen CS 77.9778 N, 14.1148 E Yes 2008 This study Arnicabreen Spitsbergen CS 78.2345 N, 16.6694 E Yes 2009 This study Bjarmebreen Spitsbergen CS 78.0136 N, 18.1316 E Yes 2011 This study Brombreen Spitsbergen CS 78.0215 N, 15.6459 E Yes 2011 This study Buckfallet Isrosa Spitsbergen CS 78.1874N, 18.6711 E Yes 2012 This study Gleditschfonna Spitsbergen CS 77.9116 N, 14.7254 E Yes 2010 This study Høgsnytbreen Spitsbergen CS 78.9828 N, 15.5752 E Yes 2011 This study Janssonbreen Spitsbergen CS 77.8965 N, 14.4408 E Yes 2008 This study Jinnbreen Spitsbergen CS 78.1580 N, 17.4489 E Yes 2009 This study Kalvdalsbreen Spitsbergen CS 77.9173 N, 15.1252 E Yes 2010 This study Kolfjellbreen Spitsbergen CS 77.8127 N, 14.9914 E Yes 2010 This study Kvitryggfonna Spitsbergen CS 78.0292 N, 17.0000 E Yes 2009 This study Königsbergbreen Spitsbergen CS 78.3455 N, 18.6477 E Yes 2012 This study Lumpbreen Spitsbergen CS 78.2147 N, 18.0001 E Yes 2012 This study Lusitaniabreen Spitsbergen CS 78.2762 N, 16.7359 E Yes 2009 This study Nord Aurdalsbre Spitsbergen CS 77.8776 N, 14.5069 E Yes 2008 This study Oppdalsbreen Spitsbergen CS 78.0913 N, 17.4344 E Yes 2011 This study Passbreen Spitsbergen CS 78.0236 N, 17.9641 E Yes 2011 This study Plogbreen Spitsbergen CS 78.0800 N, 16.1838 E Yes 2009 This study Rabotbreen Spitsbergen CS 78.2837 N, 18.0932 E Yes 2012 This study Rieperbreen Spitsbergen CS 78.1186 N, 16.1142 E Yes 2009 This study Rypefjellbreen Spitsbergen CS 77.9000 N, 15.1808 E Yes 2010 This study Sagabreen Spitsbergen CS 77.7797 N, 14.4454 E Yes 2008 This study Samuelssonbreen Spitsbergen CS 77.9229 N, 16.1613 E Yes 2011 This study Slottsbreen Spitsbergen CS 78.1000 N, 17.1358 E Yes 2009 This study Stakkbreen Spitsbergen CS 78.9891 N, 17.0427 E Yes 2009 This study Tavlebreen Spitsbergen CS 77.9797 N, 15.1509 E Yes 2011 This study Tellbreen Spitsbergen CS 78.2564 N, 16.2209 E Yes 2009 This study Tronisen Spitsbergen CS 78.1150 N, 17.0055 E Yes 2009 This study Tufsbreen Spitsbergen CS 78.0466 N, 15.6924 E Yes 2011 This study Tungebreen Spitsbergen CS 77.9621 N, 14.9511 E Yes 2010 This study Vardebreen Spitsbergen CS 78.0748 N, 13.9212 E Yes 2010 This study Vassdalsbreen Spitsbergen CS 77.8975 N, 15.0000 E Yes 2010 This study Vråbreen Spitsbergen CS 78.0640 N, 17.7310 E Yes 2011 This study Åbreen Spitsbergen CS 77.9786 N, 18.1834 E Yes 2012 This study * CS-1 Spitsbergen CS 78.1021 N, 17.2846 E Yes 2011 This study * CS-2 Spitsbergen CS 78.0878 N, 17.2759 E Yes 2011 This study * CS-3 Spitsbergen CS 78.0610 N, 16.3409 E Yes 2009 This study * CS-4 Spitsbergen CS 78.0566 N, 15.5613 E Yes 2011 This study * CS-5 Spitsbergen CS 77.9301 N, 15.0891 E Yes 2010 This study Akkarbreen Spitsbergen SS 77.1178 N, 16.5029 E Yes 2011 This study Antoniabreen Spitsbergen SS 77.4979 N, 14.9345 E Yes 2011 This study Aspelinbreen Spitsbergen SS 77.7626 N, 16.7114 E Yes 2011 This study Austre Torellbreen Spitsbergen SS 77.1731 N, 15.1419 E Yes 2011 This study Bautabreen Spitsbergen SS 76.9693 N, 16.4203 E Yes 2010 This study Belopol'skijbreen Spitsbergen SS 76.6707 N, 16.5639 E Yes 2010 This study Berrklettbreen Spitsbergen SS 77.6458 N, 15.1486 E Yes 2011 This study Bordbreen Spitsbergen SS 77.1180 N, 17.0485 E Yes 2011 This study Brattisen Spitsbergen SS 77.6053 N, 16.1581 E Yes 2011 This study Chomjakovbreen Spitsbergen SS 76.9534 N, 16.3985 E Yes 2010 This study Crammerbreane Spitsbergen SS 77.4546 N, 14.4658 E Yes 2011 This study Disbreen Spitsbergen SS 77.5338 N, 16.4424 E Yes 2011 This study Elektrobreen Spitsbergen SS 77.7755 N, 16.7168 E Yes 2011 This study Emil'janovbreen Spitsbergen SS 77.3832 N, 17.2875 E Yes 2011 This study Fiskarbreen Spitsbergen SS 77.0952 N, 16.5512 E Yes 2011 This study Frysjabreen Spitsbergen SS 77.6843 N, 15.5620 E Yes 2011 This study Goësbreen Spitsbergen SS 76.8914 N, 15.9534 E Yes 2010 This study Greenbreen Spitsbergen SS 77.6757 N, 15.7673 E Yes 2011 This study Gåsbreen Spitsbergen SS 76.9084 N, 15.9462 E Yes 2010 This study Hansbreen Spitsbergen SS 77.0486 N, 15.6380 E Yes 2011 This study Hassingerbreen Spitsbergen SS 77.4548 N, 15.7579 E Yes 2011 This study Heimbreen Spitsbergen SS 77.4861 N, 15.3577 E Yes 2011 This study Hyrnebreen Spitsbergen SS 77.0436 N, 16.2108 E Yes 2011 This study Høegh-Omdalbreen Spitsbergen SS 77.6687 N, 15.8628 E Yes 2011 This study Innifonna Spitsbergen SS 77.6630 N, 18.1762 E Yes 2012 This study Instebreen Spitsbergen SS 77.6637 N, 15.3099 E Yes 2011 This study Isingbreen Spitsbergen SS 77.1136 N, 16.6453 E Yes 2011 This study Iwobreen Spitsbergen SS 77.0933 N, 16.2366 E Yes 2011 This study Jensenbreen Spitsbergen SS 77.2671 N, 14.5261 E Yes 2011 This study Juvbreen Spitsbergen SS 77.6770 N, 16.2832 E Yes 2011 This study Kambreen Spitsbergen SS 77.0028 N, 17.1670 E Yes 2010 This study Kanebreen Spitsbergen SS 76.9782 N, 17.2169 E Yes 2010 This study Keilhaubreen Spitsbergen SS 76.6567 N, 16.9368 E Yes 2010 This study Kolkbren Spitsbergen SS 77.6676 N, 15.1724 E Yes 2011 This study Krohnbreen Spitsbergen SS 77.0741 N, 16.2424 E Yes 2011 This study Kronglebreen Spitsbergen SS 76.8242 N, 16.6841 E Yes 2010 This study Krylbreen Spitsbergen SS 77.7139 N, 15.6663 E Yes 2011 This study Kvitskarvbreen Spitsbergen SS 77.6672 N, 16.1636 E Yes 2011 This study Langkollbreen Spitsbergen SS 77.3007 N, 14.5654 E Yes 2011 This study Leinbreen Spitsbergen SS 77.4674 N, 15.3231 E Yes 2011 This study Libreen Spitsbergen SS 77.3494 N, 14.6228 E Yes 2011 This study Lundbreen Spitsbergen SS 77.6906 N, 16.1556 E Yes 2011 This study Lyngebreen Spitsbergen SS 76.6249 N, 16.6661 E Yes 2010 This study Løyndbreen Spitsbergen SS 77.7161 N, 16.0964 E Yes 2011 This study Lågryggfonna Spitsbergen SS 77.3968 N, 15.6876 E Yes 2011 This study Mathiasbreen Spitsbergen SS 76.6192 N, 16.8096 E Yes 2010 This study Matrosbreen Spitsbergen SS 77.1872 N, 17.3211 E Yes 2011 This study Mettebreen Spitsbergen SS 77.8202 N, 17.2623 E Yes 2011 This study Mikaelbreen Spitsbergen SS 77.0235 N, 16.7754 E Yes 2011 This study Morsnevbreen Spitsbergen SS 77.6382 N, 17.5714 E Yes 2011 This study Märjelenbreen Spitsbergen SS 77.4582 N, 15.4927 E Yes 2011 This study Nannbreen Spitsbergen SS 77.1354 N, 15.2816 E Yes 2011 This study Niplibreen Spitsbergen SS 77.6806 N, 15.3006 E Yes 2011 This study Nobelbreen Spitsbergen SS 77.7528 N, 16.7297 E Yes 2011 This study Odessabreen Spitsbergen SS 77.0091 N, 16.9035 E Yes 2011 This study Olsokbreen Spitsbergen SS 76.7176 N, 16.4603 E Yes 2010 This study Professorbreen Spitsbergen SS 77.0292 N, 16.8717 E Yes 2011 This study Randbreen Spitsbergen SS 76.7243 N, 17.0848 E Yes 2010 This study Raudfjellbreen Spitsbergen SS 77.2131 N, 15.0549 E Yes 2011 This study Reidbreen Spitsbergen SS 77.6207 N, 15.8472 E Yes 2011 This study Renardbreen Spitsbergen SS 77.5248 N, 14.4608 E Yes 2011 This study Revtannbreen Spitsbergen SS 77.4806 N, 15.1397 E Yes 2011 This study Ringbreen Spitsbergen SS 77.5928 N, 15.9045 E Yes 2011 This study Rokkbreen Spitsbergen SS 77.5705 N, 16.5884 E Yes 2011 This study Ryggkollbreen Spitsbergen SS 77.6608 N, 16.0459 E Yes 2011 This study Rånebreen Spitsbergen SS 77.7547 N, 16.2147 E Yes 2011 This study Saksbreen Spitsbergen SS 77.3649 N, 14.4618 E Yes 2011 This study Signybreen Spitsbergen SS 76.9661 N, 16.6736 E Yes 2010 This study Skarvisen Spitsbergen SS 77.6883 N, 15.7014 E Yes 2011 This study Skilryggbreen Spitsbergen SS 77.0850 N, 15.3946 E Yes 2011 This study Smaleggbreen Spitsbergen SS 76.9624 N, 16.4995 E Yes 2010 This study Smaubreen Spitsbergen SS 77.6561 N, 15.1541 E Yes 2011 This study Sotryggfonna Spitsbergen SS 77.4439 N, 15.9043 E Yes 2011 This study Steindolpbreen Spitsbergen SS 77.7111 N, 16.2880 E Yes 2011 This study Stepanovbreen Spitsbergen SS 77.1819 N, 17.3608 E Yes 2011 This study Suessbreen Spitsbergen SS 77.4324 N, 15.7623 E Yes 2011 This study Svalisbreen Spitsbergen SS 76.9884 N, 16.6974 E Yes 2010 This study Svartkuvbreen Spitsbergen SS 76.6765 N, 16.9399 E Yes 2010 This study Sveitssarfonna Spitsbergen SS 77.4336 N, 15.5423 E Yes 2011 This study Svingobreen Spitsbergen SS 77.8593 N, 17.0624 E Yes 2009 This study Synshovdbreen Spitsbergen SS 77.6860 N, 15.4203 E Yes 2011 This study Sysselmannbreen Spitsbergen SS 77.6185 N, 16.0496 E Yes 2011 This study Tarmbreen Spitsbergen SS 77.5528 N, 16.2809 E Yes 2011 This study Tirolarbreen Spitsbergen SS 77.4143 N, 15.7255 E Yes 2011 This study Toppbreen Spitsbergen SS 76.6358 N, 16.9046 E Yes 2010 This study Tromsøbreen Spitsbergen SS 76.9203 N, 17.0899 E Yes 2010 This study Turrsjøbreen Spitsbergen SS 77.2852 N, 14.5353 E Yes 2011 This study Tverrbreen Spitsbergen SS 77.4091 N, 14.0907 E Yes 2011 This study Tvillingbreane Spitsbergen SS 77.4345 N, 15.9563 E Yes 2011 This study Vallotbreen Spitsbergen SS 77.6365 N, 15.1698 E Yes 2011 This study Vestre Torellbreen Spitsbergen SS 77.2410 N, 14.7490 E Yes 2011 This study Vinkelbreen Spitsbergen SS 77.8099 N, 17.2694 E Yes 2011 This study Vitkovskijbreen Spitsbergen SS 76.7630 N, 16.2557 E Yes 2010 This study Werenskioldbreen Spitsbergen SS 77.0760 N, 15.2968 E Yes 2011 This study Zimmerbreen Spitsbergen SS 77.6598 N, 16.0995 E Yes 2011 This study Øydebreen Spitsbergen SS 76.8801 N, 16.7438 E Yes 2010 This study * SS-1 Spitsbergen SS 77.8563 N, 17.1765 E Yes 2009 This study * SS-2 Spitsbergen SS 77.8518 N, 18.1172 E Yes 2012 This study * SS-3 Spitsbergen SS 77.8018 N, 18.3916 E Yes 2012 This study * SS-4 Spitsbergen SS 77.7475 N, 16.9520 E Yes 2011 This study * SS-5 Spitsbergen SS 77.7238 N, 18.3158 E Yes 2012 This study * SS-6 Spitsbergen SS 77.7055 N, 18.1346 E Yes 2012 This study * SS-7 Spitsbergen SS 77.6928 N, 15.3311 E Yes 2011 This study * SS-8 Spitsbergen SS 77.6695 N, 15.9858 E Yes 2011 This study * SS-9 Spitsbergen SS 77.6305 N, 15.4804 E Yes 2011 This study * SS-10 Spitsbergen SS 77.6054 N, 15.4785 E Yes 2011 This study * SS-11 Spitsbergen SS 77.5951 N, 18.2430 E Yes 2012 This study * SS-12 Spitsbergen SS 77.5676 N, 16.3176 E Yes 2011 This study * SS-13 Spitsbergen SS 77.5582 N, 18.0639 E Yes 2012 This study * SS-14 Spitsbergen SS 77.5170 N, 17.4817 E Yes 2011 This study * SS-15 Spitsbergen SS 77.5097 N, 16.4449 E Yes 2011 This study * SS-16 Spitsbergen SS 77.4925 N, 16.7066 E Yes 2011 This study * SS-17 Spitsbergen SS 77.4727 N, 14.4167 E Yes 2011 This study * SS-18 Spitsbergen SS 77.4671 N, 16.6816 E Yes 2011 This study * SS-19 Spitsbergen SS 77.4436 N, 17.5652 E Yes 2011 This study * SS-20 Spitsbergen SS 77.4304 N, 14.6441 E Yes 2011 This study * SS-21 Spitsbergen SS 77.4273 N, 17.4473 E Yes 2011 This study * SS-22 Spitsbergen SS 77.4218 N, 16.5615 E Yes 2011 This study * SS-23 Spitsbergen SS 77.4187 N, 16.0174 E Yes 2011 This study * SS-24 Spitsbergen SS 77.4051 N, 17.2470 E Yes 2011 This study * SS-25 Spitsbergen SS 77.3666 N, 17.4962 E Yes 2011 This study * SS-26 Spitsbergen SS 77.3236 N, 17.3886 E Yes 2011 This study * SS-27 Spitsbergen SS 77.3137 N, 14.6709 E Yes 2011 This study * SS-28 Spitsbergen SS 77.1726 N, 17.2446 E Yes 2011 This study * SS-29 Spitsbergen SS 77.1268 N, 17.1259 E Yes 2011 This study * SS-30 Spitsbergen SS 77.0214 N, 16.7055 E Yes 2011 This study * SS-31 Spitsbergen SS 76.6998 N, 17.0249 E Yes 2010 This study * SS-32 Spitsbergen SS 76.6523 N, 16.5996 E Yes 2010 This study Aldousbreen Nordaustlandet NA 79.8053 N, 20.7918 E Yes 2010 This study Hårdardbreen Nordaustlandet NA 79.6950 N, 21.2342 E Yes 2010 This study Augnebreen Barentsøya BEØ 78.5325 N, 21.8003 E Yes 2010 This study Besselsbreen Barentsøya BEØ 78.5405 N, 21.5850 E Yes 2011 This study Isormen Barentsøya BEØ 78.4432 N, 22.0414 E Yes 2010 This study Willybreen Barentsøya BEØ 78.4666 N, 22.0169 E Yes 2010 This study Albrechtbreen Edgeøya BEØ 77.9308 N, 23.1531 E Yes 2010 This study Blåisen Edgeøya BEØ 78.0286 N, 21.7325 E Yes 2010 This study Gandbreen Edgeøya BEØ 77.7234 N, 22.8646 E Yes 2010 This study Hartmannbreen Edgeøya BEØ 77.3669 N, 23.0081 E Yes 2010 This study Kuhrbrenosa Edgeøya BEØ 77.5664 N, 21.9544 E Yes 2010 This study Philippibreen Edgeøya BEØ 77.7347 N, 21.8529 E Yes 2010 This study Rutenbergbreen Edgeøya BEØ 77.8904 N, 23.4136 E Yes 2010 This study Schwerdtbreen Edgeøya BEØ 77.7102 N, 21.7063 E Yes 2010 This study Seidbreen Edgeøya BEØ 77.7882 N, 22.8182 E Yes 2010 This study Skarvbreen Edgeøya BEØ 77.6066 N, 22.2214 E Yes 2010 This study Svingeldalebreen Edgeøya BEØ 78.1321 N, 21.5417 E Yes 2010 This study Sydowbreen Edgeøya BEØ 77.5741 N, 21.2093 E Yes 2010 This study Veidebreen Edgeøya BEØ 77.6632 N, 22.2904 E Yes 2010 This study Vestre Edgeøyjoøkul Edgeøya BEØ 77.8676 N, 22.7269 E Yes 2010 This study * BEO-1 Edgeøya BEØ 78.1668 N, 21.6518 E Yes 2010 This study * BEO-2 Edgeøya BEØ 78.1441 N, 21.6086 E Yes 2010 This study * BEO-3 Edgeøya BEØ 78.1427 N, 21.8600 E Yes 2010 This study * BEO-4 Edgeøya BEØ 78.1330 N, 21.9397 E Yes 2010 This study * BEO-5 Edgeøya BEØ 78.1215 N, 21.9460 E Yes 2010 This study * BEO-6 Edgeøya BEØ 78.1032 N, 21.3793 E Yes 2010 This study * BEO-7 Edgeøya BEØ 78.1022 N, 21.8918 E Yes 2010 This study * BEO-8 Edgeøya BEØ 78.0396 N, 21.7253 E Yes 2010 This study * BEO-9 Edgeøya BEØ 78.0322 N, 21.3216 E Yes 2010 This study * BEO-10 Edgeøya BEØ 77.9829 N, 22.0968 E Yes 2010 This study * BEO-11 Edgeøya BEØ 77.9139 N, 21.9787 E Yes 2010 This study * BEO-12 Edgeøya BEØ 77.8771 N, 22.5644 E Yes 2010 This study * BEO-13 Edgeøya BEØ 77.8376 N, 22.4845 E Yes 2010 This study * BEO-14 Edgeøya BEØ 77.7006 N, 21.6279 E Yes 2010 This study * BEO-15 Edgeøya BEØ 77.6989 N, 22.2611 E Yes 2010 This study * BEO-16 Edgeøya BEØ 77.6916 N, 23.0239 E Yes 2010 This study * BEO-17 Edgeøya BEØ 77.6424 N, 22.9717 E Yes 2010 This study * BEO-18 Edgeøya BEØ 77.6304 N, 23.0828 E Yes 2010 This study *BEO-19 Edgeøya BEØ 77.6227 N, 21.5578 E Yes 2010 This study * BEO-20 Edgeøya BEØ 77.5949 N, 21.4996 E Yes 2010 This study * BEO-21 Edgeøya BEØ 77.5875 N, 23.8217 E Yes 2010 This study * BEO-22 Edgeøya BEØ 77.5508 N, 23.7258 E Yes 2010 This study * BEO-23 Edgeøya BEØ 77.5084 N, 23.1169 E Yes 2010 This study * BEO-24 Edgeøya BEØ 77.4752 N, 23.1380 E Yes 2010 This study