Érosion Du Pergélisol, Transport Fluvial Et Sédimentation Marine, Côte Est De La Baie D’Hudson, Nunavik, Canada

Érosion du pergélisol, transport fluvial et sédimentation marine, côte est de la baie d’Hudson, Nunavik, Canada

Thèse

Maxime Jolivel

Doctorat en sciences géographiques Philosophiae Doctor (Ph.D.)

Québec, Canada

Résumé

Cette thèse présente une étude du système érosion-transport-sédimentation en milieu de thermokarst dans un contexte de réchauffement climatique. La zone d’étude comprend le bassin versant de la rivière Sheldrake, 5 km au nord du village Umiujaq au Nunavik, ainsi qu’une zone de 15 km2 au large de son embouchure, dans le Passage de Nastapoka, en baie d’Hudson.

Trois axes majeurs sont considérés: 1- l’étude des conditions de pergélisol et l’estimation quantitative des masses et volumes de sédiments et de carbone érodés à l’échelle du bassin versant ; 2- la mesure du régime hydrologique et sédimentaire du principal vecteur de transport, la rivière Sheldrake ; 3- la bathymétrie, la sédimentologie et la mesure des apports sédimentaires et organiques dans le milieu marin côtier au large de l’embouchure de la rivière.

Dans le bassin versant de la rivière Sheldrake, le pergélisol s’est considérablement dégradé au cours des 50 dernières années, particulièrement dans la toundra forestière. La subsidence des lithalses, des palses, des plateaux de pergélisol et des plateaux palsiques engendre la formation de mares de thermokarst. De nombreux glissements de terrain et des ravins d’érosion sont également actifs et favorisent le rejet de sédiments dans le réseau fluvial. Avec la dégradation du pergélisol, la connectivité hydrologique augmente, ce qui facilite l’évacuation des sédiments et du carbone via le cours d’eau principal.

Dans cet environnement thermokarstique, la charge sédimentaire fluviale en suspension est plus importante en été alors que les températures élevées de l’air commandent le dégel des sols, favorisant l’activation des ostioles et le déclenchement des glissements de terrain. Les pluies estivales permettent le transport et l’évacuation des sédiments en baie d’Hudson.

Parvenus en mer, les sédiments et le carbone transportés en suspension subissent une forte dispersion à cause de l’intensité des courants marins du Passage de Nastapoka. Il en résulte l’absence d’une augmentation mesurable du taux de sédimentation. En revanche, la composition isotopique du carbone sédimentaire montre que la fraction terrigène a augmenté depuis le Petit Âge Glaciaire et que ce phénomène s’est considérablement

iii accéléré vers la fin du 20ème siècle. Il est suggéré que la dégradation du pergélisol contribue à cette augmentation, quoique ce ne soit pas le seul facteur qu’on puisse invoquer.

Abstract

This thesis studies the system erosion-transport-sedimentation in a thermokastic area, in a context of warming climate. The study area encompasses the catchment of the Sheldrake River, 5 km north of the village Umiujaq, Nunavik, and a 15 km2 area off its mouth, in the Nastapoka Sound, in Hudson Bay.

Three main axes are considered: 1- study of permafrost conditions and quantitative estimate of the volumes and masses of eroded sediment and organic carbon at the scale of the catchment; 2- measurements of the hydrological and sedimentary regime of the main vector of transport, the Sheldrake River; 3- bathymetry, sedimentology and measurements of mineral and organic inputs in the coastal marine environment, off the river mouth.

In the Sheldrake River catchment, permafrost has considerably degraded during the last 50 years, particularly in the forested tundra. Subsidence of lithalsas, palsas, permafrost plateaus and peat plateaus leads to the formation of thermokarst ponds. Many landslides and erosion gullies are also active and favor inputs of sediments in the fluvial network. Because of permafrost decay, hydrological connectivity increases, facilitating evacuation of sediment and carbon through the river.

In this thermokarstic environment, the fluvial sedimentary load in suspension is more important during summer when high air temperatures provoke soils thawing, favoring frostboils activation and triggering of landslides. Summer rainfalls allow sediment transport and evacuation in Hudson Bay.

Once in the sea, the sediments and carbon in suspension are dispersed because of the intensity of the marine currents in the Nastapoka Sounds. This results in an absence of a measurable increase of sedimentation rates. However, the isotopic composition of sedimentary carbon shows that the terrestrial fraction has increased since the Little Ice Age and that this trend has significantly accelerated since the end of the 20th century. It is suggested that permafrost decay contributes to this increase, although it is not the only proposed source.

Table des matières

Résumé ...... iii

Abstract ...... v

Table des matières ...... vii

Liste des tableaux ...... xi

Liste des figures ...... xiii

Remerciements ...... xv

Avant-propos ...... xvii

CHAPITRE 1 - Introduction générale ...... 1

1.1 Problématique générale ...... 1

1.2 Objectifs généraux ...... 2

1.3 Objectifs spécifiques ...... 3

1.4 Région d’étude ...... 4

1.5 Méthodologie ...... 10 1.5.1 Cartographie de la distribution et de la dégradation du pergélisol et estimation quantitative des masses et volumes de sédiments et de carbone érodés ...... 10 1.5.2 Mesures du régime hydrologique et sédimentaire de la rivière Sheldrake ...... 11 1.5.3 Bathymétrie, sédimentologie et mesure des apports sédimentaires et organiques dans le Passage de Nastapoka au large de l’embouchure de la rivière Sheldrake ...... 11

1.6 Structure de la thèse ...... 12 1.6.1 Chapitre 2 ...... 12 1.6.2 Chapitre 3 ...... 12 1.6.3 Chapitre 4 ...... 13 1.7 Références ...... 13

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CHAPITRE 2 - Thermokarst and export of sediment and organic carbon in the Sheldrake River watershed, Nunavik, Canada...... 17

2.1 Résumé ...... 17

2.2 Abstract ...... 18

2.3 Introduction ...... 18

2.4 Study area and regional setting...... 20

2.5 Permafrost degradation processes ...... 25

2.6 Methodology ...... 28 2.6.1 Permafrost and thermokarst mapping ...... 28 2.6.2 Landslides identification and sediment and organic matter released ...... 31 2.6.3 Estimation of organic carbon and sediment fluxes from connected thermokarst ponds ...... 32 2.6.4 Hydrological connectivity ...... 36

2.7 Results ...... 36 2.7.1 Spatiotemporal evolution of permafrost ...... 36 2.7.2 Activity of landslides, active layer failures and gullies between 1957 and 2009 ...... 41 2.7.3 Hydrological connectivity ...... 41 2.7.4 Potential TSS, DOC and OM flow in the system from CTPs ...... 42

2.8 Discussion ...... 43 2.8.1 Toward a permafrost-free landscape ...... 43 2.8.2 Importance of the activity of landslides, active layer failures and gullies ...... 47 2.8.3 Increase of sediments and OM released by landslides, active layer failures and gullies ...... 48 2.8.4 Impact of an increasing connectivity on organic carbon and mineral sediment yields ...... 49

2.9 Conclusion ...... 50

2.10 Acknowledgements ...... 51

2.11 References ...... 51

CHAPITRE 3 - Hydrological regime and sediment transport in a river flowing in a thermokarst landscape, Sheldrake River, Nunavik, Quebec...... 59

3.1 Résumé ...... 59

3.2 Abstract ...... 59 viii

3.3 Introduction ...... 60

3.4 Study area ...... 62

3.5 Methods ...... 65 3.5.1 Field instrumentation and laboratory analyses ...... 65 3.5.2 Stage/discharge calibration ...... 66 3.5.3 NTU/TSS calibration ...... 67 3.5.4 Active layer depth ...... 68

3.6 Results ...... 68 3.6.1 Interannual and seasonal variations of discharge ...... 68 3.6.2 Turbidity regime over the thawing season ...... 70 3.6.2.1 Turbidity variations during the snowmelt period ...... 70 3.6.2.2 Relation between turbidity and summer air temperature (2010 and 2013) ...... 72 3.6.3 Total suspended sediment export estimates ...... 73

3.7 Interpretation and discussion ...... 74 3.7.1 Hydrological regime of a typical subarctic river ...... 74 3.7.2 Impact of active layer thawing and thermokarst processes on sediment annual regime ...... 75 3.7.3 Yearly measured rates of transport vs assessment of mass lost by thermokarst ...... 80

3.8 Conclusion ...... 81

3.9 Acknowledgements ...... 82

3.10 References ...... 83

CHAPITRE 4 - Morphostratigraphy and recent sedimentation in Nastapoka Sound, Eastern Coast of Hudson Bay ...... 89

4.1 Résumé ...... 89

4.2 Abstract ...... 89

4.3 Introduction ...... 90

4.4 Study area ...... 92

4.5 Methods ...... 95 4.5.1 Bathymetry and recovery of cores ...... 95 4.5.2 Laboratory methods ...... 97 4.5.3 Dating ...... 98 4.5.4 C, N, δ13C and δ15N analysis ...... 99 4.5.6 Historical changes in the fraction of terrestrial organic matter ...... 99

4.6 Results ...... 100 4.6.1 Sea floor morphology and distribution of surface sediments ...... 100 4.6.2 Chronology ...... 101 4.6.3 Lithology and grain size ...... 103 4.6.4 Elemental and isotopic composition of sedimentary OM ...... 107

4.7 Discussion ...... 108 4.7.1 Sedimentation rates and chronology ...... 108 4.7.2 A complex sedimentary regime ...... 110 4.7.2.1 A dynamic area ...... 110 4.7.2.2 Disturbed sediment deposits: core 05 ...... 111 4.7.2.3 Bottom current deposits ...... 113 4.7.3 Recent sedimentation: increase of terrigenous influence ...... 113 4.7.3.1 δ13C and C/N molar ratio ...... 114 4.7.3.2 δ15N ...... 115 4.7.3.3 Historical changes in terrigenous OM contribution since the LIA ...... 117

4.8 Conclusion ...... 121

4.9 Acknowledgments ...... 121

4.10 References ...... 122

Chapitre 5 - Conclusion générale ...... 129

Liste des tableaux

Table 1.1 Répartition des dépôts meubles en fonction de la zone biogéographique ...... 8

Table 2.1 Surficial deposits distribution in the Sheldrake River catchment...... 23

Table 2.2 Lithalsas, palsas, total permafrost coverage and changes between 1957 and 2009. a CTTP is connected thermokarst pond from palsa collapse...... 37

Table 2.3 Thermokarst pond coverage, connected thermokarst pond coverage and changes between 1957 and 2009. a Thermokarst pond resulting from palsa collapse...... 38

Table 2.4 Number and average size of lithalsas, palsas and thermokarst ponds in the forest tundra area, and changes between 1957 and 2009...... 38

Table 2.5 Characteristics and measurements of the landslides, the active layer detachment failures and the gullies of the Sheldrake River catchment in 2009.a shrub tundra;b forest tundra ...... 40

Table 2.6 Number, volume and mass changes in landslides and gullies between 1957 and 2009...... 41

Table 2.7 Yields of sediment in suspension and dissolved organic carbon in tons from connected thermokarst ponds to the drainage system, in 2009 and 2010...... 43

Table 2.8 Permafrost and thermokarst pond changes at different sites in northern Québec from references...... 44

Table 3.1 Hydrological features of the Sheldrake River in 2009, 2010, 2011, 2012 and 2013 during the ice free period...... 70

Table 3.2 Comparison between the summer periods of turbidity in 2010 and in 2013...... 73

Table 3.3 Sediment exported by the Sheldrake River and specific values in 2010 and 2013. *In 2013, data acquisition stopped on 19 October...... 74

Table 3.4 Sediment released by thermokarst in the total fluvial exports...... 81

Table 4.1 Features of the 25 sediment cores extracted in April 2009. Sediment textures and structures are given from visual inspection and CT scan imagery, except cores 3, 5, 10, 13, 23 and 24 (grain size analyses). IRD is ice-rafted debris. (a) and (b) in sedimentary structures referred to the given (a) and (b) sediment texture...... 97

Table 4.2 14C dates ...... 103

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Liste des figures

Figure 1.1 Carte de localisation de la zone d'étude ...... 5

Figure 1.2 Courbe du relèvement isostatique postglaciaire de l'est de la baie d'Hudson (Lavoie et al., 2012)...... 6

Figure 1.3 Carte de distribution du pergélisol au Québec (Allard et al., 2012)...... 9

Figure 2.1 Location of the Sheldrake River catchment showing the position of the tree line and the maximum elevation reached by the Tyrell Sea (marine limit)...... 21

Figure 2.2 Distribution of surficial deposits on the Sheldrake river catchment. White areas are bedrock outcrops...... 23

Figure 2.3 Change in the maximum active layer depth in the VDT lithalsa between 2001 and 2009. Vertical axis shows the depth in cm below the ground surface...... 25

Figure 2.4 Permafrost decay features in the Sheldrake River catchment. Photographs: (a) typical thermokarst landscape in the forest tundra area; (b) a frostboil evolving into an active layer failure; (c) palsa peat blocks falling in a pond (photo: D. Sarrazin); (d) active layer failure detachment on the edge of a high permafrost plateau; (e) network of eroding gullies between permafrost plateaus; (f) active layer slide occurring in spring 2010; note the frostboils on the top of the plateau (photo: D. Sarrazin)...... 27

Figure 2.5 Map of the degradation of the permafrost in percentage and location of landslides, active layer detachment failures and gullies. The catchment was divided into 41 numbered land systems units; segmentation was based on the visual criteria of permafrost percentage cover, topographic style and type of surficial deposits...... 31

Figure 2.6 Method for estimating missing volume from landslides along river banks: satellite image, field photograph and geometrical calculation...... 31

Figure 2.7 Hygrographs (water level) of the Sheldrake River from mid-June to mid-October in 2009 and in 2010. Every rain event with an impact on the river discharge is numbered...... 34

Figure 2.8 Change in drainage density caused by lithalsas and palsas decay in land system 24. .... 42

Figure 2.9 Gradient of percentage of permafrost degradation between 1957 and 2009 from the Hudson Bay landward, with position of the tree line. Each point corresponds to one land system. Sandy areas are excluded...... 45

Figure 3.1 (A) Location of the study area; (B) The Sheldrake River catchment with the distribution of surficial deposits and the organization of the drainage network. The rest of the area is bedrock outcrops...... 63

Figure 3.2 (a) Relationship between water level (m) and discharge (m3/s); (b) relationship between turbidity (NTU) and total suspended sediment (TSS) (g/m3)...... 68

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Figure 3.3 Hydrographs for 2009, 2010, 2011, 2012 and 2013. Curve of discharge starts when discharge reaches 12 m3/s, which is supposed to be the first signs of ice break up...... 69

Figure 3.4 Hydrological and turbidity regime of the Sheldrake River depending on air temperature and rainfall events in 2010: rain (mm/day), air temperature (°C), discharge (m3/s) and turbidity (NTU). Data are measured on a daily basis...... 71

Figure 3.5 Hydrological and turbidity regime of the Sheldrake River depending on air temperature and rainfall events in 2013: rain (mm/day), air temperature (°C), discharge (m3/s) and turbidity (NTU). Data are measured on a daily basis...... 72

Figure 3.6 Annual air temperature in Umiujaq between 1998 and 2013. No data are available in 1999, 2001 and 2002...... 76

Figure 3.7 Turbidity (UTN), thaw front depth (cm), rain (mm/day) and cumulative degrees-day ≥ 0⁰C from 1 May to 1 October 2010...... 77

Figure 3.8 Turbidity (UTN), rain (mm/day) and cumulative degrees-day ≥ 0⁰C from 1 May to 1 October 2013...... 78

Figure 3.9 (a): active and flowing frostboil on the slope of a lithalsa (photo: Denis Sarrazin); (b): Typical thermokarst ponds, the strong turbidity is caused by frostboils activity and surface runoff; (c): typical landslide on a permafrost bank along the Sheldrake River; (d): small delta at the confluence between a thermokarst gully and the Sheldrake river...... 79

Figure 4.1 Bathymetric map and location of the coring sites. Bathymetric data south of the unmapped areas come from Girard Thomas (2009)...... 93

Figure 4.2 Erosional landforms on the sea floor. (A) Trailing-spits (encircled); (B) erosion channel (arrow)...... 101

210 210 210 Figure 4.3 Profiles of Pb total and Ln Pb in excess (ln PbXs) VS depth in the sea bed, and concentration of 137Cs. Gray zones indicate the surface mixed layer. The dashed lines show the supported 210Pb...... 103

Figure 4.4 CT-scan image, mean grain size (µm), magnetic susceptibility (k) and CT-number of cores 05, 10 and 24. A zoom of the CT-scan image of core 05 highlights the convolutes and the erosion contact. Light gray bands show erosion contacts in cores 05 and 10; darker gray bands highlight sand beds in core 24. The black star indicates 1900 AD inferred from 210Pb profiles. .... 105

Figure 4.5 Physical, elemental and isotopic profiles of cores 03, 13 and 23. The black star indicates 1900 AD inferred from 210Pb profiles...... 106

Figure 4.6 Average grain size frequency for core 13...... 107

Figure 4.7 δ13C vs C/N in the Nastapoka Sound (white squares) compared with values in Lac Guillaume Delisle (black triangles) and in the rest of Hudson Bay (black circle)...... 115

Figure 4.8 Variation in the proportion of terrestrial organic matter in cores 03, 13 and 23 since 1750 AD. Arrows highlight common trends...... 117

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Remerciements

De nombreuses personnes, et quelques personnages, ont rendu cette aventure universitaire et personnelle possible. Évidemment, ma gratitude la plus sincère va à Monsieur Michel Allard, mon directeur de recherche. Parce qu’à partir d’un simple mail posté fin 2006 sur la toile par un individu en recherche d’emploi, il a su me faire confiance, m’accorder le temps, les moyens et le financement nécessaires pour mener à bien mes recherches. Il m’a octroyé une réelle liberté d’expression scientifique allant du choix du sujet de thèse, aux méthodes d’analyse employées en passant par la possibilité de me rendre sur le terrain à sept reprises. Je lui suis également reconnaissant pour m’avoir permis d’enseigner à de nombreuses reprises, une expérience très enrichissante, captivante et valorisante.

Je remercie les membres de mon comité de thèse, Patrick Lajeunesse et Guillaume St-Onge pour nos fructueuses et stimulantes discussions ainsi que pour l’utilisation de matériel comme le banc MSCL du GEOTOP de Rimouski. Merci également à Mr Daniel Fortier qui a accepté d’être l’examinateur externe de cette thèse.

Je suis reconnaissant envers Mr Bill Doidge qui nous a permis d’utiliser son bateau et son matériel pour la campagne de terrain de l’été 2008.

Je remercie les personnes qui ont participé aux analyses relatives aux carottes de sédiment et ont su répondre à mes nombreuses questions : Guillaume Labrecque du Laboratoire de radiochronologie de l’Université Laval, Jacques Labrie du Geotop de Rimouski, Stéphane Prémont et Louis Frédéric Daigle de l’Institut National de la Recherche Scientifique, Wendy C. Abdi et Paul Middlestead du G.G. Hatch Laboratories.

Au sein du Centre d’études nordiques, quatre âmes charitables, souriantes et disponibles (parfois taquines) ont été d’une aide inestimable : Emmanuel l’Hérault, Denis Sarrazin, Carl Barrette, Mickael Lemay. Plusieurs autres personnes m’ont apporté leur connaissance ou leur assistance, à un moment ou à un autre de ces 6 années : Maud Audet Morin, Donald Cayer, Marc-André Ducharme, Catherine Falardeau-Marcoux, René-Charles Bernier, Valérie Maton Dufour, Jonathan Roger, Luc Cournoyer, Alexandre Normandeau et bien d’autres qui me pardonneront leur omission.

Cette thèse a été financée par le Conseil de Recherches en Sciences Naturelles et en Génie du Canada (CRSNG), le réseau Arcticnet, le programme ADAPT et le Centre d’études nordiques.

Je remercie la communauté inuit d’Umiujaq. J’y ai rencontré des personnes accueillantes et chaleureuses. J’ai vécu dans ce village une expérience mémorable au côté d’un peuple fort et authentique qui je l’espère saura s’adapter aux nouveaux défis qui l’attend.

Merci à mes amis d’avoir été des amis, d’ici ou d’ailleurs, et d’avoir fait ce que je suis aujourd’hui. Des chemins se sont séparés, de nouveaux se sont ouverts et la vie continue.

Finalement, mes pensées vont vers ma famille qui a su respecter mes choix et accepter la distance et mon absence. Il est parfois difficile de vivre loin des siens, mais si la vie offre sans cesse de nouvelles pistes à suivre, je suis convaincu d’avoir pris la bonne.

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Avant-propos

Cette thèse présente une synthèse de mes travaux de recherche de doctorat, réalisé sous la direction du professeur Michel Allard. Les chapitres 2, 3 et 4 sont rédigés sous la forme d’articles scientifiques. Ce choix implique une redondance dans la description de la zone d’étude. Cependant, chaque chapitre peut ainsi être lu indépendamment des autres. Le chapitre 2 est publié dans la revue à comité de lecture Journal of Geophysical Research (Jolivel et Allard, 2013). Les chapitres 3 et 4 le seront prochainement. Ces articles sont rédigés en anglais, langue internationale, afin de communiquer le plus efficacement possible les résultats à la communauté scientifique. Je suis l’auteur principal des chapitres et des figures de cette thèse. J’ai été responsable des travaux de terrain et de la majorité des analyses en laboratoire. Pour faciliter la lecture, les références de chaque chapitre sont répertoriées à la fin de chacun d’entre eux. Les 3 chapitres-articles de cette thèse sont une suite. Par conséquent, dans les chapitres 3 et 4, il est parfois fait références aux chapitres 2 et 3.

Chapitre 1 – Introduction générale

Chapitre 2 - Thermokarst and export of sediment and organic carbon in the Sheldrake River watershed, Nunavik, Canada. M. Jolivel et M. Allard

Chapitre 3 - Hydrological regime and sediment and organic carbon transport in a river flowing in a thermokarst landscape, Sheldrake River, Nunavik, Quebec. M. Jolivel et M. Allard

Chapitre 4 - Morphostratigraphy and recent sedimentation in Nastapoka Sound, Eastern Coast of Hudson Bay. M. Jolivel, M. Allard et G. St-Onge

Chapitre 5 – Conclusion générale

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CHAPITRE 1 - Introduction générale

1.1 Problématique générale

Le pergélisol est un phénomène climatique. Sa température, inférieure à 0°C pendant au moins deux années consécutives, est en étroite relation avec les températures atmosphériques et donc, avec le climat. Dans le contexte actuel de réchauffement climatique global, sa dégradation représente un enjeu majeur tant d’un point de vue environnemental que social, économique et politique.

Le pergélisol des régions arctiques et subarctiques renferme 1700 Pg de carbone, soit ~ 50% du carbone terrestre global ou le double du volume de carbone atmosphérique mondial (Schuur et al., 2008, Tarnocai et al., 2009; MacDougall et al., 2012). Le dégel de sols gelés depuis des centaines ou des milliers d’années va permettre la dégradation de la matière organique désormais disponible à l’action bactérienne, engendrant la formation de gaz à effet de serre (C02 et CH4), eux-mêmes contributeurs du changement climatique. C’est le principe de la rétroaction positive. De plus, si les prévisions d’augmentation des températures de l’air poursuivent la tendance actuelle, les processus thermokarstiques vont s’intensifier et entrainer de nombreux bouleversements du relief tels que l’érosion et les mouvements de versants. Le carbone organique et les sédiments ainsi libérés dans les écosystèmes fluviaux, littoraux et marins auront des impacts significatifs sur la production primaire et conséquemment, sur les chaines alimentaires. Il devient donc urgent d’acquérir de solides connaissances sur les processus de dégradation du pergélisol et le devenir des produits de cette dégradation.

Au Nunavik, sur la côte est de la baie d’Hudson, plusieurs études spatio-temporelles ont démontré que le pergélisol discontinu présent dans les dépôts meubles de la Mer de Tyrrell s’est dégradé d’environ 40 % depuis les 50 dernières années (Laberge et Payette, 1995; Payette et al., 2004; Marchildon, 2007; Vallée et Payette, 2007; Fortier et Aubé-Maurice, 2008). Cette dégradation se caractérise notamment par la subsidence des buttes (lithalses et palses) et plateaux de pergélisol et la formation de mares de thermokarst. Le paysage se transforme littéralement sous nos yeux : un environnement sec, bien drainé et couvert d’une

1 végétation de toundra laisse place à un paysage humide, aquatique, avec une végétation arbustive et forestière de plus en plus dense. La multiplication des mares de thermokarst provoque une réorganisation des réseaux naturels de drainage. L’augmentation du taux d’érosion du pergélisol, à travers par exemple les glissements de terrain, engendre une augmentation de la charge sédimentaire transportée par les cours d’eau.

Dans la région d’étude, le Centre d’études nordiques (CEN) a mené de nombreuses recherches sur le pergélisol depuis les années 1980. De précieuses connaissances ont été acquises sur la formation historique du pergélisol (Allard et Séguin, 1987; Calmels et al., 2008), sa structure interne générale (Allard et Rousseau, 1999), le contenu et le type de glace (Calmels et Allard, 2004, 2008), et les processus et les facteurs de formation et de dégradation des lithalses et des palses (Calmels et al., 2007, 2008; Delisle et al., 2003; Larouche, 2010). Nous présentons ici la suite logique de ces recherches : que devient le matériel sédimentaire qui constituait auparavant ces formes de pergélisol?

L’originalité de ce travail tient d’abord dans sa continuité logique. Schématiquement, on se propose en effet de suivre les sédiments depuis leur érosion thermokarstique sur le bassin versant jusqu’à leur dépôt dans un bassin sédimentaire marin. Par conséquent, chacun des trois chapitres principaux de cette thèse se concentrera sur un environnement particulier : périglaciaire (érosion), fluvial (transport) et marin (sédimentation).

Cette thèse utilise une approche novatrice et systémique dans le but d’améliorer la compréhension de la dégradation du pergélisol. Par son innovation en termes de zone d’étude, de méthodes et de résultats, ce travail de recherche s’avère d’une contribution scientifique significative et appelle de nouveaux travaux afin de préciser les processus originaux mis en évidence à travers ces trois chapitres et mieux comprendre leur impact sur les écosystèmes.

1.2 Objectifs généraux

L’objectif général de cette thèse est de mieux comprendre le fonctionnement du système érosion-transport-sédimentation en contexte thermokarstique. Pour cela, il est nécessaire de décrire les processus géomorphologiques propres à la dégradation du pergélisol, quantifier

les volumes et les masses de sédiments et de carbone produits par le thermokarst, mesurer le transport sédimentaire dans le réseau fluvial et analyser la façon dont les sédiments et le carbone apportés à la mer sont intégrés dans la dynamique sédimentaire marine. Par l’étude méticuleuse d’un bassin versant subissant un thermokarst actif, nous avons estimé les volumes de sédiments et de carbone organique érodés, mesuré les masses impliquées dans le transport fluvial et, finalement, échantillonné et analysé les lieux et processus de déposition marine de ce « nouveau » matériel disponible. L’approche systémique utilisée permet une exploration minutieuse des domaines de la géomorphologie périglaciaire, fluviale et marine en milieu subarctique.

1.3 Objectifs spécifiques

Les objectifs spécifiques sont les suivants :

1. Quantifier la dégradation récente (1957-2009) du pergélisol à l’échelle d’un bassin versant. 2. Décrire et interpréter les changements géomorphologiques liés à la dégradation du pergélisol entre 1957 et 2009. 3. Estimer le volume potentiel de sédiment et de carbone organique libéré dans le réseau de drainage par les formes d'érosion thermokarstique entre 1957 et 2009. 4. Quantifier l'impact de l'augmentation du nombre de mares de thermokarst et des changements d'organisation du réseau hydrologique, entre 1957 et 2009, en termes de volumes de sédiment et de carbone organique remobilisés dans le réseau de drainage. 5. Documenter le régime hydrologique et sédimentaire d’une rivière drainant un bassin versant subissant un thermokarst actif. 6. Établir un lien entre le transport sédimentaire et les processus thermokarstiques. 7. Décrire la dynamique morpho-sédimentaire dans le Passage de Nastapoka. 8. Caractériser les apports d’origine terrigène dans la sédimentation marine récente du Passage de Nastapoka.

1.4 Région d’étude

Les recherches terrestres ont été menées dans le bassin versant de la rivière Sheldrake (76 km2) qui se jette dans la baie d’Hudson à ~5 km au nord du village inuit d’Umiujaq (56°37’N; 76°32’W). La volet « marin » a été réalisé dans le Passage de Nastapoka, situé entre l’Archipel des Iles Nastapoka et la côte du Québec. L’échantillonnage se concentre sur une zone de 3 km sur 5 km, directement au large de l’embouchure de la rivière Sheldrake (Figure 1.1).

Figure 1.1 Carte de localisation de la zone d'étude

La région d’étude est composée de deux ensembles géologiques. L’arrière-pays repose sur des formations archéennes de granito-gneiss appartenant à la province tectonique du Supérieur. Les reliefs sont typiques du Bouclier canadien, soit une pénéplaine ondulée, d’une altitude d’environ 250 m, entaillée de vallées encaissées d’orientation générale est- ouest. La zone côtière, le Passage de Nastapoka ainsi que les Iles Nastapoka sont constitués de roches volcano-sédimentaires protérozoïques de la province de Churchill (groupe de Nastapoka). La base de cette formation affleure au front des cuestas des Iles Nastapoka et des crêtes asymétriques immergées du Passage de Nastapoka. A environ 6 km du rivage s’étire une rangée de collines parallèles au trait de côte et d’une altitude atteignant 400 m. Finalement, une bande côtière descend de ces reliefs pour rejoindre la baie d’Hudson avec des pentes inférieures à 3%.

Dans la région, la déglaciation wisconsinienne a débuté vers 8200 ans étal. BP (Lavoie et al., 2012). Le front glaciaire s’est d’abord stabilisé sur les reliefs côtiers, menant à la mise en place, en condition sous-marine, de la ceinture de drift de Nastapoka (Lajeunesse et Allard, 2003a et 2003b). La côte, déprimée par le poids de l’inlandsis, fut envahie par la mer de Tyrrell. La vitesse du relèvement isostatique atteignait alors ~1 m/an. Depuis lors, celui-ci s’est ralenti pour atteindre actuellement une vitesse de 1.3 m/siècle (Lavoie et al., 2012) (Figure 1.2).

Figure 1.2 Courbe du relèvement isostatique postglaciaire de l'est de la baie d'Hudson (Lavoie et al., 2012).

Par conséquent, les reliefs de la région d’étude ont été fortement modelés par le passage de l’Inlandsis laurentidien et l’invasion de la mer de Tyrrell. A l’est du lac Sheldrake, la limite marine atteint 220 m ; elle s’abaisse à 204 m, perpendiculairement à la côte, dans le prolongement de la vallée de la rivière Nastapoka (Allard et Séguin, 1985 ; Lavoie, 2000). Hormis les hauts reliefs, le bassin versant de la rivière Sheldrake fût donc totalement submergé par la mer de Tyrrell, comblant en partie les vallées de limons argileux en leur conférant des fonds plats. A l’intérieur des terres, les dépôts de sable et de gravier associés aux deltas fluvioglaciaires marquent l’altitude maximale atteinte par la mer de Tyrrell. Sur le versant des collines côtières, les dépôts grossiers des éventails sous-marins de contact glaciaire ont été remobilisés sous formes de terrasses et de plages lors de la baisse du niveau marin relatif associée au relèvement isostatique postglaciaire. Suite à l’émersion des terres, la végétation et les sols se sont développés. Les fonds plats et mal drainés des vallées ont favorisé la formation de tourbières.

Typiquement, la stratigraphie des fonds de vallée du bassin versant de la rivière Sheldrake est la suivante: le socle rocheux du bouclier canadien, du till, du limon argileux glaciomarin 6

d'eau profonde, des sables d'exondation mis en place lors du retrait des eaux de la mer de Tyrrell, un horizon tourbeux.

La région est aujourd’hui soumise à l’influence d’un climat subarctique (moyenne annuelle des températures entre 1960 et 2010 : -5.3°C) dont la variabilité est étroitement liée à la présence de la baie d’Hudson, prolongement méridional de l’Océan arctique. Lorsque la baie est libre de glace, de juin à décembre, le climat est frais (10°C en août) et humide, favorisant la formation de brouillards persistants. L’hiver, la présence de l’immensité glacée de la baie d’Hudson engendre un climat sec, venteux et froid (-24°C en janvier). Les précipitations moyennes annuelles sont de l’ordre de 530 mm, réparties inégalement au cours de l’année (60% de juillet à novembre), 40% tombant sous forme de neige. Dans la région d’Umiujaq, l’indice de gel a été évalué à 3056 degrés-jours et celui de dégel à 1014 degrés-jours (Boyd, 1973). Dans la région d'étude, les rivières sont généralement englacées au début du mois de novembre et la débâcle intervient en moyenne dans la première moitié du mois de mai. Par la suite, le niveau d'eau reste élevé pendant la fonte des neiges puis décroit progressivement pour atteindre son niveau estival à la fin du mois de juin ou au début du mois de juillet (Déry et al., 2005; Chapitre 3).

L’influence climatique de la baie d’Hudson commande la distribution de la végétation. La zone côtière appartient au domaine de la toundra arbustive composée principalement de lichens, d’herbacées, d’éricacées, de krummholz d’épinette noire et de divers arbustes comme le bouleau nain (Betula glandulosa), le saule (Salix planifolia) et l’aulne (Alnus crispa) (Payette et Rochefort, 2001). A une dizaine de kilomètres du littoral, la toundra arbustive cède progressivement la place à la toundra forestière dominée par les mêmes espèces précédemment nommées en plus de peuplements d’épinettes noires et de mélèzes (Larix laricina) dans les vallées abritées (Payette et Rochefort, 2001). La transition entre ces deux entités biogéographiques, la limite des arbres, coupe littéralement le bassin versant de la rivière Sheldrake en deux parties. La toundra arbustive est présente sur 43% du bassin versant, tandis que la toundra forestière occupe le territoire restant (Tableau 1.1).

Dépôts meubles Dont (%)

km2 % km2 % Limon Sable Tourbe Toundra forestière 43 57 13 54 84 11 5 Toundra arbustive 33 43 11 46 88 12 0 Total 76 100 24 100

Table 1.1 Répartition des dépôts meubles en fonction de la zone biogéographique

Finalement, ce gradient thermique provoqué par la présence de la baie d’Hudson influence la distribution et la dégradation du pergélisol qui s’est installé dans la région lors des périodes froides de l’Holocène récent (Allard et Seguin, 1987). La région étudiée fait partie de la zone de pergélisol discontinu, caractérisée par une concentration importante de palses, de plateaux palsiques, de lithalses et de plateaux de pergélisol présents dans les tourbières et les dépôts fins de la mer de Tyrrell (Figure 1.1). En 2009, le pergélisol est ainsi présent dans 20% des dépôts meubles. Ces formes surplombent de quelques mètres les terrains bas environnants et les tourbières. Elles renferment des lentilles de glace de ségrégation qui constituent jusqu’à 60% de leur volume interne. La profondeur du pergélisol dans les dépôts de surface est en moyenne inférieure à 20 m, sa température moyenne se situe entre -0.5 et -2°C. L'épaisseur de la couche active varie de 1 à 3 m en fonction du type de substrat et la profondeur d'amplitude thermique annuelle nulle est généralement inférieure à 4 m (Calmels et al., 2007, 2008; Calmels et Allard, 2008; Fortier et al., 2008).

Alors que les palses et les plateaux palsiques, surmontés d’un couvert de tourbe, sont surtout présents dans la toundra forestière, les lithalses et les plateaux de pergélisol abondent dans la toundra arbustive, où le pergélisol est deux fois plus épais (Lévesque et al., 1988). Ces formes de terrain sont les principaux signes de la présence de pergélisol dans la région, auxquels il faut ajouter les ostioles et dans une moindre mesure, les sols polygonaux. Si la présence de pergélisol est avérée dans le socle rocheux, en particulier dans les affleurements côtiers, sa distribution régionale dans le roc reste inconnue. En revanche, il a été grossièrement estimé qu’environ 50% des dépôts meubles du bassin versant de la rivière Sheldrake était sous régime pergélisolé (Lévesque et al., 1988).

Dans la région, le thermokarst, qui désigne les formes et processus associés à la subsidence d’un pergélisol riche en glace, est très actif. Il se remarque par la présence de nombreuses mares de thermokarst, de coulées de gélifluxion et de formes d’érosion liées à l’approfondissement de la couche active du pergélisol. On estime que l’activité thermokarstique a débuté à la fin du Petit Âge Glaciaire, puis s’est accélérée considérablement en réponse au réchauffement climatique récent (Payette et al., 2004).

Figure 1.3 Carte de distribution du pergélisol au Québec (Allard et al., 2012).

1.5 Méthodologie

La méthodologie suivie durant cette thèse est adaptée aux trois volets de la recherche : 1- cartographie de la distribution et de la dégradation du pergélisol et estimation quantitative des masses et volumes de sédiments et de carbone érodés, 2- mesures du régime hydrologique et sédimentaire du principal vecteur de transport, la rivière Sheldrake; 3- bathymétrie, sédimentologie et mesure des apports sédimentaires et organiques dans le milieu marin côtier au large de l’embouchure de la rivière. Les données brutes ont été recueillies au cours de six campagnes de terrain.

1.5.1 Cartographie de la distribution et de la dégradation du pergélisol et estimation quantitative des masses et volumes de sédiments et de carbone érodés

Le bassin versant de la rivière Sheldrake fût délimité grâce à un modèle numérique d’élévation créé à partir d’une carte topographique à l’échelle 1 : 50 000 avec une résolution altitudinale de 20 m (Ressources Naturelles, Faune et Parcs Québec). Les formes et les processus dynamiques de pergélisol sur le territoire du bassin ont été cartographiés dans le logiciel ArcGis. Afin d’établir un état des lieux récent de la distribution du pergélisol, les formes de pergélisol et de thermokarst furent numérisée sur une image satellitaire GeoEye de 2009 (résolution : 0.6 m). Les flux de carbone et de matières en suspension issus de l’augmentation de la connectivité hydrologique ont été estimés en sélectionnant les mares de thermokarst connectées au réseau de drainage. Un travail similaire fût réalisé sur des photographies aériennes de 1957 (résolution : ~ 1 m). La soustraction des deux cartes ainsi créées (distribution du pergélisol et des mares de thermokarst en 2009 et en 1957) a permis de créer une carte de dégradation du pergélisol. Des observations de terrain, des photographies et une vidéo pris lors de survols en hélicoptère ont permis de réduire les incertitudes. Afin d’estimer les volumes érodés, plusieurs formes d’érosion ont été mesurées directement sur le terrain. Leur totalité fût ensuite répertoriée sur l’image satellitaire et les photographies aériennes, puis reportée sur les cartes qui ont donné lieu à la comparaison et à la soustraction quantitative des volumes disparus.

1.5.2 Mesures du régime hydrologique et sédimentaire de la rivière Sheldrake

La méthodologie repose en majeure partie sur les données recueillies grâce à l’installation d’une station de jaugeage à 2 km de l’embouchure de la rivière Sheldrake. Deux capteurs y mesurent en continu le niveau d’eau (Levellogger) et la turbidité (OBS 3+). La vitesse du courant fût mesurée grâce à un courantomètre 2100-LX et la surface mouillée à l’aide d’un échosondeur et d’une simple perche. Des prélèvements d’eau ont été réalisés afin de connaitre les concentrations de sédiment en suspension et de carbone organique à des niveaux d’eau différents. Les données de niveau d’eau furent converties en débit (m3/s) et les données de turbidité en taux de matières en suspension (g/m3). Les analyses de concentrations en matière en suspension, en carbone organique total et en carbone organique dissous furent réalisées au Laboratoire Environnex et au laboratoire de l’Institut National de la Recherche Scientifique (INRS), tous deux situés à Québec.

1.5.3 Bathymétrie, sédimentologie et mesure des apports sédimentaires et organiques dans le Passage de Nastapoka au large de l’embouchure de la rivière Sheldrake

Les relevés bathymétriques du Passage de Nastapoka ont été réalisés à l’été 2008, à bord du bateau MV Katherine-Anne, opéré par Bill Doidge, à l’aide d’un écho-sondeur Raymarine DSM300 couplé d’un sonar à balayage latéral Edgetech 4100P. Vingt-cinq carottes ont été prélevées en avril 2009, à partir du couvert de glace, à l’aide d’un carottier à gravité K-B ® (16 kg). Toutes les carottes entières ont été passées au scanner médical (CT-Scan) de l’INRS à Québec, pour en imager les structures sédimentaires. Par la suite, elles ont été coupées en deux dans le sens de la longueur, photographiées, décrites visuellement puis analysées sur le banc GEOTEK MSCL (Multi Sensor Core Logger) de l’Institut des Sciences de la Mer de Rimouski.

Six carottes ont été retenues selon leur localisation géomorphologique et leur potentiel de datation de l’histoire sédimentaire récente: dans les bassins sédimentaires profonds et le long d’un transect perpendiculaire à l’embouchure de la rivière Sheldrake. Des analyses granulométriques ont été réalisées au laboratoire de géomorphologie de l’université Laval. Des datations au 14C, 210Pb et 137Cs ont été confiées au laboratoire de radiochronologie du

Centre d’études nordiques. Finalement, trois carottes ont été choisies et échantillonnées à des fins d’analyses élémentaires (C, N) et isotopiques (δ13C and δ15N), aux G.G. Hatch Isotope Laboratories de l’université d’Ottawa.

1.6 Structure de la thèse

Le cœur de cette thèse est composé de trois articles scientifiques (chapitres 2, 3 et 4). Le premier est déjà publié, les deux autres seront soumis prochainement. La présente introduction constitue le premier chapitre alors qu’une synthèse des résultats et une conclusion générale forment le 5ème et dernier chapitre.

1.6.1 Chapitre 2

Le 2ème chapitre de cette thèse est intitulé “Thermokarst and export of sediment and organic carbon in the Sheldrake River watershed, Nunavik, Canada”. Sous la forme de cartes et de statistiques exhaustives, ce chapitre présente d’abord un état des lieux de la distribution et de la dégradation du pergélisol à l’échelle du bassin versant de la rivière Sheldrake. Par la suite, on propose une estimation des volumes et des masses de sédiment et de carbone organique issus du thermokarst et potentiellement remobilisés dans le réseau de drainage. Deux types de changements géomorphologiques sont considérés : les glissements de terrain et la surface couverte par les mares de thermokarst connectées au réseau de drainage (connectivité hydrologique). L’objectif de ce chapitre est de confirmer que le pergélisol se dégrade et de cerner l’échelle spatiotemporelle du processus. Conséquemment, des sédiments sont intégrés dans le réseau fluvial. Ceci constitue l’hypothèse de départ de la thèse. Ce chapitre est publié dans la revue Journal of Geophysical Research Earth Surface (Jolivel et Allard, 2013).

1.6.2 Chapitre 3

Le 3ème chapitre se nomme “Hydrological regime and sediment transport in a river flowing in a thermokarst landscape, Sheldrake River, Nunavik, Quebec”. Ce chapitre présente des hydrogrammes et des valeurs de turbidité de la rivière Sheldrake afin de caractériser son

régime hydrologique et sédimentaire. Un évènement de forte turbidité, relié aux processus périglaciaires et thermokarstiques, est également décrit et discuté. L’objectif de ce chapitre est d’apporter de nouvelles connaissances sur la réponse des cours d’eau aux variations climatiques dans un contexte de dégradation du pergélisol. Ce manuscrit, qui est une suite logique au chapitre 2 sera prochainement soumis pour publication à une revue sceintifique.

1.6.3 Chapitre 4

Le 4ème chapitre présenté est intitulé “Morphostratigraphy and recent sedimentation in Nastapoka Sound, Eastern Coast of Hudson Bay”. Ce travail présente une carte bathymétrique ainsi que des images de sonar à balayage latéral du Passage de Nastapoka, au large de l’embouchure de la rivière Sheldrake. On y trouve également une étude physique, géochimique et isotopique de six carottes de sédiments pour lesquelles on a également calculé les taux de sédimentation et évalué les sources d’apport en carbone. L’objectif de ce chapitre est d’une part, de préciser les processus de sédimentation et d’érosion subaquatiques dans le Passage de Nastapoka, et d’autre part de mettre en évidence des changements récents dans la sédimentation d’origine terrigène et de les interpréter dans un contexte de changements environnementaux, particulièrement depuis le Petit Âge Glaciaire. Ce chapitre sera soumis à la revue avec comité de lecture Estuarine, Coastal and Shelf Science.

1.7 Références

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Calmels, F., & Allard, M. (2008). Segregated ice structures in various heaved permafrost landforms through CT Scan. Earth Surface Processes and Landforms, 33, 209–225. Allard, M., Lemay, M., Barrette, C., L’Hérault, E., Sarrazin, D., Bell, T. & Doré, G. (2012). Permafrost and climate change in Nunavik and Nunatsiavut: Importance for municipal and transportation infrastructures. In: Allard, M. and M. Lemay (Eds), Nunavik and Nunatsiavut: From science to policy. An Integrated Regional Impact Study (IRIS) of climate change and modernization. ArcticNet Inc., Quebec City, Canada, p. 171-197. Boyd, D.W., 1973. Normal freezing and thawing degree-days for Canada, 1931-1960. Canada. Department of Environment, Atmospheric Environment Service, Publ. CL 1 4-73: 38.

Calmels, F., Allard, M., & Delisle, G. (2007). Development and decay of a lithalsa in Northern Québec: A geomorphological history,Geomorphology. Geomorphology, 97(3-4), 287–299.

Calmels, F., Delisle, G., & Allard, M. (2008). Internal structure and the thermal and hydrological regime of a typical lithalsa: significiance for permafrost growth and decay. Canadian Journal of Earth Sciences, 45, 31–43.

Delisle, G., Allard, M., Fortier, R., Calmels, F., & Larrivée, É. (2003). Umiujaq, Northern Québec: innovative techniques to monitor the decay of a lithalsa in response to climate change. Permafrost and Periglacial Processes, 14, 375–385.

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CHAPITRE 2 - Thermokarst and export of sediment and organic carbon in the Sheldrake River watershed, Nunavik, Canada.

2.1 Résumé

Une analyse spatiotemporelle de la dégradation du pergélisol entre 1957 et 2009, validée par des mesures sur le terrain, a été réalisée sur un bassin versant de 76 km2 afin d’estimer les volumes de sédiments et de carbone organique libérés par l’activité thermokarstique. La zone d’étude se situe à la limite des arbres, dans la zone de pergélisol discontinu, sur la côte est de la baie d’Hudson. Les lithalses et les palses sont les formes de pergélisol les plus répandues alors que les mares de thermokarst, les glissements de terrain et les ravins d’érosion sont les principaux signes de dégradation du pergélisol. Les résultats montrent que 21% du pergélisol existant en 1957 a disparu en 2009, entrainant une hausse de 96% de la surface couverte par les mares de thermokarst et une augmentation de 46 à 217% du nombre de formes d’érosion actives. Une augmentation de la connectivité hydrologique associée à la dégradation du pergélisol a potentiellement engendré une multiplication par 1.6 des volumes de sédiment et de carbone organique rejetés dans le réseau de drainage. L’activité des glissements de terrain et des ravins a crû de 12 à 38%, contribuant également à l’augmentation des volumes de sédiments et de carbone organique libérés dans le système fluvial. Des différences significatives dans la dégradation du pergélisol et dans les rejets de sédiment et de carbone organique ont été observées le long d’un gradient ouest-est, entre des sites proches de la côte de la baie d’Hudson, dans la toundra arbustive, et des sites localisés à l’intérieur des terres au sein de la toundra forestière. Les mares de thermokarst dans la toundra forestière rejettent dans le réseau fluvial 2.3 fois plus de sédiment et de carbone organique dissous par unité de surface que dans la toundra arbustive. Toutefois, malgré cette intense activité thermokarstique, le bassin versant de la rivière Sheldrake ne semble pas rejeter proportionnellement plus de sédiments et de carbone organique qu’un bassin versant sans pergélisol.

2.2 Abstract

A spatiotemporal computation of permafrost decay covering the period from 1957 to 2009 and validated by field investigations, was made over a 76 km2 river catchment straddling the tree line, in the discontinuous permafrost zone, east of Hudson Bay, in order to estimate the amounts of sediments and organic carbon released by thermokarst. Lithalsas and palsas are the dominant permafrost landforms whereas thermokarst ponds, landslides, active layer failures and gullies are the main features of permafrost degradation. Results show that 21% of the existing permafrost in 1957 had disappeared in 2009, resulting in a 96% growth of the thermokarst pond cover and a 46 to 217% increase of the number of active erosion landforms. An increase of stream connectivity related with the degradation of permafrost potentially allowed for an increase of sediments and carbon delivery to the main stream by a factor of 1.6. Volume of active landslides and gullies also increased by 12 to 38%, enhancing sediment and organic matter yields. Significant differences in permafrost degradation and in sediment and carbon inputs were observed along a east-west transect, from sites located at the head of the watershed near the tree line, and sites located downstream close to the Hudson Bay coast. Thermokarst ponds in the forest tundra area released 2.3 times more sediments and dissolved organic carbon per unit of area in the fluvial system than in the shrub tundra area. Despite these yields by thermokarst, the Sheldrake River catchment currently does not seem to be yielding proportionally more sediments and carbon than a permafrost free river catchment.

2.3 Introduction

Recent climate changes have induced geomorphic perturbations in arctic and subarctic landscapes because of severe terrain disturbances caused by the thawing of ice-rich permafrost (Jorgenson et al., 2001; Jones et al., 2011). These impacts are particularly conspicuous in the discontinuous permafrost zone as permafrost extent is decreasing rapidly, resulting in an increase in number and extent of thermokarst lakes in areas where soil drainage is poor, such as clays and peatlands (Laberge and Payette, 1995; Matthews et al., 1997; Camill, 1999; Osterkamp et al., 2000; Beilman et al., 2001; Payette et al., 2004; Vallée and Payette, 2007). However, in spite of measured spatial reduction of permafrost

areas, very little field data exist on the volumes and masses of mineral and organic sediment released to stream systems by permafrost decay.

Most papers have focused on the spatial decay of permafrost at the scale of palsa bogs (Laprise and Payette, 1987; Laberge and Payette, 1995; Zuidhoff and Kolstrup, 2000; Vallée and Payette, 2007) and at the regional scale (Vitt et al., 2000a; Thibault and Payette, 2010; Sannel and Kuhry, 2011). However, the impact of permafrost thawing on sediment release and contribution to solid discharge in river catchments is still poorly known. Some attempts were made to measure sediment and carbon outputs from stream basins in degrading permafrost regions (Kokelj et al., 2005; Bowden et al., 2008; Olefeldt and Roulet, 2012). Both the original and the transitional basin morphology during permafrost terrain degradation are likely to exert some control on the paths and discharge of thermokarst products, particularly carbon. Such an assessment is needed as the organic matter (OM) that was stored in the permafrost is recycled in thermokarst ponds as particulate organic matter (POM) and dissolved organic matter (DOM), which are involved in the biogeochemical production of methane (CH4) and dioxide gas (CO2). Furthermore, a fraction of the released fine-grained sediments and carbon also flows into the fluvial system and, ultimately, to the sea where it likely contributes to changes into the marine food web and the sedimentary system.

The main objective of this paper is to provide a better understanding of the processes of geomorphic change in a river catchment where permafrost is being impacted by thermokarst. The study investigates the release of sediments and carbon from permafrost decay in the fluvial system with emphasis on the potential export of fine sediments (silty clay) as suspension load, dissolved organic carbon (DOC) and particulate organic carbon (POC). We describe erosion and fluvial sediment transport within a catchment whose morphology, particularly stream connectivity, is being transformed by thermokarst. Our approach consists of estimating sediment and carbon volume and mass changes resulting from recent and current thermokarst and related denudation processes within the watershed. We also tested the hypothesis that the increase in number and extent of thermokarst ponds has increased the overall sediments and carbon load of the fluvial system studied.

In order to obtain this new catchment-scale knowledge and to better understand regional mass transfers related to permafrost thawing, the specific objectives of the study were (1) to quantify permafrost decay between 1957 and 2009 in a river basin located in the discontinuous permafrost zone, (2) to quantify the area and size of new thermokarst ponds and erosion features since 1957, and (3) to assess the amounts and fractions of organic matter, organic carbon and mineral sediments sourcing from recent permafrost degradation that are mobilized in the reorganized watershed.

2.4 Study area and regional setting

The Sheldrake River is 25 km long from its source in Sheldrake Lake to the shore of Hudson Bay. Inland, it flows over glacially-eroded Precambrian granitic gneiss in a hilly (200-250 m a.s.l.) landscape partially covered by late glacial and postglacial sediments. The six last kilometres of the river run in a valley carved across Late Proterozoic sedimentary- volcanic rocks which form a chain of coastal hills (400 m a.s.l.) parallel to the arc-shaped coastline of the eastern Hudson Bay. For the final two kilometres, the river flows on a gentle structural coastal slope (<3°) following the dip angle of a thick basalt layer. It flows into Hudson Bay 8 km north of the Inuit village of Umiujaq (56°37’N ; 76°32’W) (Figure 2.1).

Following the receding front of the Laurentide Ice Sheet eastward and inland, between 8000 and 7400 cal. yr BP (Lajeunesse and Allard, 2003a; Lajeunesse, 2008; Lavoie et al., 2012), the Tyrrell Sea inundated the whole Sheldrake River watershed (Figure 2.1). The studied valley and associated tributary valleys were filled with postglacial marine silty clay, reaching up to 85 m in thickness (Allard and Seguin, 1985). Therefore, silty clay accounts for 86% of surface deposits (Table 2.1). Sandy and gravelly ice-contact deltas (Lajeunesse and Allard, 2003a) mark the marine limit at an elevation of 220 m a.s.l. at the east end of the Sheldrake River basin (Allard and Seguin, 1985). Near the mouth of the river, sand and gravel deposits anchored on the Nastapoka Hills form a fan-like apron sloping from the hills to the actual sea-shore (Allard and Seguin, 1985; Lajeunesse and Allard, 2003a, 2003b). In flat and poorly-drained inland valleys, peat covers silt and sand deposits with an

average thickness of one meter (Lévesque et al., 1988) (Figure 2.2). Surficial deposits cover 32% of the watershed. Bedrock and water (lakes) make up the balance.

Figure 2.1 Location of the Sheldrake River catchment showing the position of the tree line and the maximum elevation reached by the Tyrell Sea (marine limit).

Figure 2.2 Distribution of surficial deposits on the Sheldrake river catchment. White areas are bedrock outcrops.

Area Total surficial Silty clay Sand Peat deposits

2 2 km km % % %

Shrub tundra 33 11 88 12 0 Forest tundra 43 13 84 11 5 Total 76 24 86 11 3

Table 2.1 Surficial deposits distribution in the Sheldrake River catchment.

The climate is subarctic, with cold winters (-24°C in January) and cool summers (10°C in August). Since 1990, mean annual air temperatures increased from -5.2°C to -3.9°C; 40% of the 550 mm average annual precipitation falls as snow (Environment Canada, 2010). The cooling impact of Hudson Bay generates a west-east climate gradient from the shoreline inward, resulting in the tree line being parallel to the coastline, about 15 km inland (details on vegetation can be found in Payette and Rochefort, 2001 and Bhiry et al., 2011). The tree line refers to the first occurrence of isolated trees in the landscape (Payette, 1983). Here, compared to regional estimates of the tree line position (Payette, 1983), it has been slightly moved to west to include valleys with scattered spruces.

The 76 km2 studied watershed is located at the boundary between the isolated and sporadic permafrost zone (2 to 50% areal cover) and the widespread discontinuous permafrost zone (50 to 90%) (Allard and Seguin, 1987a). According to Lévesque et al. (1988), in 1957, permafrost covered approximatively 50% of surficial deposits in the Sheldrake River basin. The rest of the surficial deposits are permafrost free and consist of taliks, forested and shrubby hollows and wetlands. Lithalsas (mineral permafrost mounds), permafrost plateaus (elongated and wide mineral permafrost landforms), palsas (peaty permafrost mounds or permafrost mounds with a peat cover) and peat plateaus (elongated and wide peaty permafrost landforms) are the only permafrost landforms (for simplicity, the term “lithalsa” is used here to include permafrost plateaus and the term “palsa” encompasses peat

23 plateaus). All these heaved landforms stand above the surrounding terrain by several meters (generally 3 to 5 m) due to the development of ice segregation lenses formed by cryosuction of soil water in freezing fine-grained, frost susceptible soils (Pissart, 1985, 2002). Their volumetric ice content varies between 50 and 80% (Calmels and Allard, 2004; Calmels et al., 2008). Permafrost thickness in the frozen mounds and plateaus varies typically from 10 to 15 m (Lévesque et al., 1988). The active layer is ~1.5 m deep on clay soils (Delisle et al., 2003; Calmels and Allard, 2004, 2008; Calmels et al., 2007) and is currently thickening (Figure 2.3). It is about 0.6 m in peat on tops of palsas (Marchildon, 2007). One of our thermistor cables attests of the presence of permafrost in bedrock near the coast in the village of Umiujaq. However, regional distribution of permafrost in bedrock is unknown.

In northern Québec, palsas and lithalsas formed during cold periods of the Late Holocene, i.e. 1500 to 1000 BP, and during the Little Ice Age (LIA) (Couillard and Payette, 1985; Allard and Seguin, 1987a, 1987b; Lavoie and Payette, 1995; Payette and Delwaide, 2000; Arlen-Pouliot and Bhiry, 2005; Marchildon, 2007). Since the end of the LIA, numerous permafrost landforms have degraded due to global warming and to increased snow precipitations (Seguin and Allard, 1984; Vitt et al., 2000a; Payette et al., 2004; Arlen- Pouillot and Bhiry, 2005). Permafrost mounds at different stages of degradation are found within the region (Calmels et al., 2007). The main geomorphological impact associated with this decay is an increasing number of thermokarst ponds and erosion on hill slopes (Allard et al., 1987).

The Sheldrake River watershed was selected for this study because several previous studies in the region provide baseline information on permafrost spatial distribution, landforms, permafrost thickness, ground ice content and thermal regime. This strategic area straddles the tree line, in a region of intensive degradation of permafrost, and is representative of the eastern coast of Hudson Bay landscape. Meteorological stations and sites for monitoring permafrost decay, such as the BGR lithalsa (informal name for Bundesanstalt für Geowissenschaften und Rohstoffe, a German research center which collaborated with the Centre d’études nordiques in the early 2000’s) and the VDT lithalsa (for Vallée des Trois, informal name given to a valley east of the village of Umiujaq), provide quantitative data

that help to understand landscape adaptation to permafrost decay (see Figure 2.1 for locations). Finally, a water level gauge installed in 2008 2 km upstream from the Sheldrake River mouth provides useful hydrological information applicable to the catchment.

Figure 2.3 Change in the maximum active layer depth in the VDT lithalsa between 2001 and 2009. Vertical axis shows the depth in cm below the ground surface.

2.5 Permafrost degradation processes

In the studied watershed, the first evidence of degradation is the observed recent thawing of permafrost mounds (Figure 2.4a). Surface and slopes of lithalsas are affected by thaw settlement, gelifluction, active layer failure and slopewash erosion. Alteration of the thermal balance of the ice-rich mounds first creates a summit depression which deepens and enlarges summer after summer, retaining water and forming a thermokarst pond surrounded by an ice-poor ridge (Calmels et al., 2007). Clayey sediments are brought to the expanding pond by overland flow on the bare surface of frostboils, which keep bringing fine-grained sediment to the surface through overturning (Figure 2.4b). During rain events, clay is eroded and channelled in furrows between the frostboils into the adjacent ponds or

25 directly into the stream network (Seguin and Allard, 1984). Small active layer slides along the shore of the ponds often release sediments as well (Allard et al., 1987).

On palsas and peat plateaus, extension cracks inherited from the time of heave occur on the swelled surface of the mound and divide the peat cover into blocks that collapse in surrounding ponds (Figure 2.4c), streams and rivers, as frequently observed in North America (Allard and Seguin, 1987a; Allard et al., 1987; Lévesque et al., 1988; Vitt et al., 1994; Laberge and Payette, 1995; Osterkamp and Romanovsky, 1999; Beilman et al., 2001; Payette et al., 2004; Arlen-Pouillot and Bhiry, 2005) and elsewhere (Matthews et al., 1997; Zuidhoff and Kolstrup, 2000). In peat plateaus, ground subsidence can initiate the thermokarst pond which expands through peat block erosion (Sannel and Kuhry, 2011). Some palsas also gradually lose their peat cover through erosion, particularly by the wind (Seppälä, 2003). When the peat cover is totally eroded, the underlying silt is exposed at the ground surface. Since silt has a higher thermal conductivity than peat, the active layer gets thicker. Frostboils become active.

During the subsidence of palsas and lithalsas, slope retreat delivers peat and sediments in the expanding ponds. Therefore, the thermokarst ponds act as sediment traps and are very turbid (Breton et al., 2009). A pond isolated from the fluvial drainage network is a small sedimentation basin within a closed catchment (Bouchard et al., 2011). Consolidation of postglacial silty clay underlying thermokarst ponds makes the soil nearly impermeable once the permafrost is completely thawed. Consequently, vertical water infiltration is negligible. Drainage of the pond only occurs when it overflows or when it becomes connected to the main drainage network through some newly eroded creek channels. Some mounds however are directly located along river banks. When they decay, the morphology of the river channel is affected (Vallée and Payette, 2007) and their sediments are then directly taken in charge as river load.

Landslides occur along steep river banks, particularly in areas of permafrost plateaus, in the western part of the studied watershed. Small active layer failures occur on the steep upper edge of permafrost plateaus (Figure 2.4d). They are active during summer thaw (Seguin and Allard, 1984) and lead to the failure of “a thin veneer of vegetation and mineral soil and subsequent movement over a planar inclined surface” (McRoberts and Morgenstern,

1974). Finally, gully erosion (Figure 2.4e) occurs between adjacent permafrost plateaus and can create networks of gullies in some permafrost fields. They are directly connected to the drainage network through elongated ravines.

2.6 Methodology

The methodology is based on precise mapping, estimates of landforms changes in volume and mass, observations of flow regime in the catchment and observations of active processes.

2.6.1 Permafrost and thermokarst mapping

The Sheldrake River watershed was first delineated using a digital elevation model generated from a topographic map (contour lines: 20 m, Ressources Naturelles, Faune et Parcs Québec). We choose to exclude the Sheldrake Lake catchment as it has no visible signs of erosion, being surrounded by wetlands and bedrock outcrops. Only the area drained by the Sheldrake River and its tributaries was accounted for in measurements of erosion and mass computation. Surficial deposits were mapped using 1957 aerial photographs (1:40,000) (~1 m resolution), a 2009 GeoEye satellite image (0.6 m resolution), field observations, and revisions of previous maps by Lévesque et al. (1988) at a smaller scale (1:50,000). The 1957 aerial photographs were geo-referenced with the 2009 satellite image using ~ 20 control points per photographs distributed throughout the area. Control points mainly consisted of rock outcrops.

Every permafrost mound and thermokarst pond in the catchment was delineated and mapped in the ArcGis software for both 1957 and 2009. This method allowed for a much higher precision than previous studies which used area estimates (Lévesque et al., 1988). For 1957, the use of a stereoscope with enlarging lenses (3X) allowed for the detection of typical heaved permafrost landforms. For 2009, permafrost features and thermokarst ponds were precisely mapped on the high resolution satellite image. Lithalsas are pitted with light active frostboils. Palsas have a more uniform surface and the brownish colour of the peat cover is characteristic. The whitish thermokarst ponds are often surrounded by rim ridges and look like sinkholes in the landscape. Field observations, photographs and videos taken from helicopter flights over the whole basin area helped to verify uncertain geological contacts and validate mapping. Not all thermokarst ponds are associated with a rim ridge, particularly in wetlands, in peat bogs and in fen areas. Coalescence of several ponds leads

to the creation of larger lakes and unstructured wet areas. In this study, every water body resulting from permafrost decay was considered as a thermokarst pond. The margin of error is estimated to less than 5% for the delineation of each pond and permafrost mound. Moreover, no other aerial photographs were available between 1957 and 2009. Thus, possible variations in the rate of change were not considered between these two dates.

In order to detect variability in changes within the catchment and to assess the relative contribution in sediment and carbon fluxes from different sectors in the watershed, the catchment area was classified into 41 numbered units of land systems. Segmentation was based on the visual criteria of permafrost percentage cover, topographic style and type of surficial deposits (Figure 2.2). Two permafrost maps were created and subtracted to produce a map of permafrost degradation between 1957 and 2009 (Figure 2.6).

The presence of permafrost in sandy areas (land systems 1, 13, 35, 36, 40; for location, see Figure 2.5) is difficult to confirm because no typical periglacial landform indicates presence or absence of permafrost below the surface. Those areas are mostly wind-eroded surfaces supporting small dune ridges and were considered as permafrost areas without vegetation as suggested by Lévesque et al. (1988), whom made numerous geo-electrical surveys on these terrain types. No visible signs of thermokarst have been observed in these areas and sand deposits represent only 11% of total catchment in unconsolidated deposits. The study focused on the potential export of fine sediments and carbon as suspension load in the fluvial network. Therefore, sand deposits are shown on the map, but were not included in the computations.

2.6.2 Landslides identification and sediment and organic matter released

The yields of sediment and carbon from decaying permafrost were estimated through geometrical measurements on air photographs supported by ground checks and measurements at typical sites. The measured landforms included landslides, gullies and small active layer failures.

Three landslides were visited in the field and seven were identified on the satellite image. They consist of non vegetated slopes in thawed and collapsed sediments along steep river banks that are constantly destabilized by fluvial toe erosion. Geometrical parameters (height, length, horizontal distance and breath) were measured in the field and on the satellite image. Thereafter, the eroded volume in each landslide was estimated with trigonometrical functions (Figure 2.6). Because of the high rate of the isostatic uplift in this region (Lavoie et al., 2012), fluvial incision is still in progress and maintains steep river banks, particularly along large permafrost plateaus. Consequently, we assume that the original land surface was a cube volume without any slope.

Figure 2.6 Method for estimating missing volume from landslides along river banks: satellite image, field photograph and geometrical calculation.

Slope of three active layer failures and one system of gullies were measured in the field and respectively six and 19 such landforms were identified on the satellite image in the basin.

In order to estimate volumes of eroded sediments, active erosion surfaces (i.e. bare and unstable) were measured on the satellite image. An average scar depth of 0.75 m with a slope of 40° for active layer failures and an average scar depth of 0.9 m with a slope of 45° for gullies between two plateaus were estimated by field measurements. Erosion landforms less than 3 m in length were not considered in this study.

Volume was converted into mass by assuming a porosity of 0.35 and a density of the solid particles of 2.7 t/m3, which are typical values for the glaciomarine clays (Holtz and Kovacks, 2009). The volume to mass conversion for landslides takes into account a 1.5 m thick active layer without ice content and an underlying permafrost with an average ice content of 60% (Calmels et al., 2007, 2008; Calmels and Allard, 2008; Fortier et al., 2008). For active layer failures and gullies, the ice content is considered null since erosion only takes place in the active layer. We assume that the totality of the volume of missing sediment is released in the fluvial system and that OM is exported in DOM or POM.

A large complex of gullies is active in a scrubland sector in the center of the catchment. The abundance of shrubs since at least 1957, which retain a thick snow cover, indicates that this sector has no permafrost. This was confirmed by electrical resistivity surveys conducted by Lévesque et al. (1988). Although morphometric changes were measured in this sector as well, this area was excluded from the computation of permafrost-induced sediment quantities since it does not actually contribute to thermokarst-induced sediment input into the river system.

2.6.3 Estimation of organic carbon and sediment fluxes from connected thermokarst ponds

All ponds connected to the drainage network were inspected to assess the volume of turbid water that is evacuated when they overflow during rain events. Connected thermokarst ponds (CTP) in 2009 were first delineated using the high resolution GeoEye image. Sometimes, draining streams were not visible, but wet linear shrubby hollows connecting ponds to the river drainage system act as flow paths during precipitation events. This is confirmed by the presence of small deltas at the confluence with permanent streams. To

assess the 1957 situation, every pond that was connected in 2009 and already present on the old air photographs was considered as CTP.

To calculate the coverage of new connected thermokarst ponds (NCTP), i.e. the area of CTP created between 1957 and 2009 (area NCTP), we used the following equation:

area NCTP = area CTP09 – area CTP57 (1) where area CTP57 and area CTP09 are the areas covered (m2) by CTP in 1957 and in 2009. To assess the current release of total suspended sediment (TSS) and DOC from CTP, it is assumed that the volume of water (with known contents of DOC and TSS) overflowing into the drainage network is equal to the volume formed by area CTP09 and the thickness of a layer of cumulated rain that generates overflow. This volume must be considered as a minimum because the area of the drainage basin of each thermokarst pond was not taken in account.

The hydrographs from the gauging station (Figure 2.7) show that every rain event >2 mm/day has an impact on the Sheldrake River water level. Therefore, during a rain event, we can assume that flow occurs through CTPs. The study focused on the period from mid- June, at the end of the freshet, to mid-October, just before freeze back. A rain event is defined here as a period of rain of one or several consecutive days leading to a peak of discharge, i.e. a water stage on the hydrograph (Figure 2.7).

Figure 2.7 Hygrographs (water level) of the Sheldrake River from mid-June to mid- October in 2009 and in 2010. Every rain event with an impact on the river discharge is numbered.

The volume of water released by NCTP in the drainage network (W) between the 15 June and the 15 October is summarized by the expression:

W = area NCTP × P (2) where P is the layer of cumulated rain (m) falling on the catchment between June 15th and October 15th in 2009 and 2010.

Finally, exports of TSS and DOC from NCTP were calculated as:

expDOC = area NCTP × P × avTSS× 10-6 (3)

expTSS = area NCTP × P × avDOC× 10-6 (4) where expDOC and expTSS are the mass of DOC and TSS (tons) released by NCTP between June 15th and October 15th of each year; avTSS and avDOC are respectively the average concentration of TSS and DOC (g.m-3) in NCTP. 10-6 is the conversion factor from grams to tons.

Estimating avTSS and avDOC in the catchment is difficult given that the turbidity of ponds and streams is highly variable in space and time (~4400 ponds in 2009). We used concentration values from Breton et al. (2009). These authors had selected in our study area 16 thermokarst ponds they had estimated representative of the variety of colors (related to sediment and organic matter concentrations) and development phases among numerous thaw ponds in the region. They have measured an average concentration of 19.8 g.m-3 of TSS (16 ponds) and 4.9 g.m-3 of DOC (5 ponds). The TSS is principally composed of silt and clay (Breton et al., 2009). They also provided measurements of TSS and DOC concentrations in 17 other thermokarst ponds in the wider region of eastern Hudson Bay coast, which all fall in the same range as those taken in our study area. In the Sheldrake river catchment, concentrations of DOC range from 1.3 to 11.5 g.m-3 (median: 4.3; σ: 3.05; standard error of the mean: 0.76) and concentrations of TSS range from 5.3 to 39.7 g.m-3 (median: 15.5; σ: 12.9; standard error of the mean: 1.13) (Breton et al., 2009). From these data, we simply used average concentration values of 20 g.m-3 of TSS and 5 g.m-3 of DOC for the total population of thermokarst ponds in our study area. This calculation assumes a linear relationship between precipitation and DOC/TSS loading, even if a strong flow dependency of export can exist (Finlay et al., 2006). It also ignores effects of connectors, which, if vegetated, can filter out TSS and influence DOC concentration. Evaporation between rain events is not considered. Estimates are very likely to change with more detailed sampling of concentration in the future. Despite large uncertainties, this approach provides an acceptable order of magnitude though still rough and preliminary.

2.6.4 Hydrological connectivity

Providing a precise overview of the first order drainage network of the whole catchment is complicated because of the discontinuous extent of surface sediments between rock outcrops and the fragmented nature of the landscape. In order to test and illustrate how thermokarst may have affected stream connectivity, a ~ 0.2 km2 flat sub-basin located in the forest tundra area (land system 24) was selected. This sub-basin was selected because of the presence of palsas, the rate of permafrost decay and the increase in thermokarst pond coverage, which are similar to the mean value of the forest tundra area. The stream density (Dd) is defined as the cumulated length of water tracks (L) divided by the area of the sub- basin (A) (Horton, 1945) (Equation 5).

Dd = Σ L / A (5)

2.7 Results

2.7.1 Spatiotemporal evolution of permafrost

The mapping of the entire basin of the Sheldrake River provides an overview of permafrost conditions on the eastern coast of Hudson Bay. In 2009, permafrost in the Sheldrake River watershed was present in 20% of surficial deposits, while it was 26% in 1957. In 2009, permafrost was present in 32% of surficial deposits in the shrub tundra area, while it was present in only 10% of the surficial deposits in the forest tundra (Table 2.2).

Lithalsas Palsas Total permafrost m2 % m2 % m2 %

Shrub tundra

1957 3,324,779 34.3 1042 0.1 3,335,199 34.4 2009 3,097,963 32 9052 0.1 3,107,015 32 Change -226,816 -6.8 -1368 -13.1 -228,184 -6.8 Forest tundra

1957 1,860,067 16.1 243,008 0.9 2,103,075 18.3 2009 1,082,923 9.4 103,873 2.1 1,186,796 10.3 Change -777,144 -42 -139,135 -57.3 -916,279 -43.4 Total

1957 5,184,846 24.4 253,428 1.2 5,438,274 25.6 2009 4,180,886 19.7 112,925 0.5 4,293,811 20.2 Change -1,003,960 -19 -140,503 -55.4 -1,144,463 -21

Table 2.2 Lithalsas, palsas, total permafrost coverage and changes between 1957 and 2009. a CTTP is connected thermokarst pond from palsa collapse.

Complete statistics of permafrost decay and thermokarst ponds between 1957 and 2009 for the 41 land systems are available in appendixes 1 and 2. Synthesised data are presented in Table 2.2 and Table 2.3.

Total permafrost coverage decreased by 21% from 1957 to 2009. However, the degradation is very unevenly distributed in space, with an average loss of permafrost of only 7% in the shrub tundra area, compared to a permafrost loss averaging 43% in the forest tundra, east of the tree line (Figure 2.5).

In 1957, thermokarst ponds covered 0.56 km2. In 2009, this value has doubled to reach 1.09 km2 (+ 96%). In the forest tundra, the number of ponds has increased from 2,958 ponds in 1957 to 4,442 in 2009 (+50%) (Table 2.4). Logically, thawing ponds replace permafrost mounds and the trend follows the rate of permafrost degradation.

Thermokarst ponds Connected thermokarst ponds CTPPa

2 2 2 m % m % m Shrub tundra

1957 121,457 1.3 25,664 21 \ 2009 210,629 2.2 44,437 21 913 Change 89,172 73 18,773 73 \

Forest tundra

1957 435,319 3.8 126,969 29 \ 2009 882,712 7.7 197,067 22 20,597 Change 447,393 103 70,098 55 \

Total

1957 556,776 2.6 152,633 27 \ 2009 1,093,341 5.2 241,504 22 21,510 Change 536,565 96 88,871 58 \

Table 2.3 Thermokarst pond coverage, connected thermokarst pond coverage and changes between 1957 and 2009. a Thermokarst pond resulting from palsa collapse.

Lithalsas Palsas Thermokarst ponds Number Size (m2) Number Size (m2) Number Size (m2) 1957 1062 1806 256 954 2958 154 2009 1100 1274 116 895 4442 205 Change 38 -532 -140 -59 1484 51 Change (%) 4 -29 -55 -6 50 33

Table 2.4 Number and average size of lithalsas, palsas and thermokarst ponds in the forest tundra area, and changes between 1957 and 2009.

Flat floored poorly-drained inland valleys were the most affected areas. Permafrost degradation was generally above 50% and the surface covered by thermokarst ponds increased by 100% or more since 1957 (see land systems 19, 27, 33, 34, 41 in Appendixes 1 and 2). In the four areas where palsas and lithalsas occur together, i.e. in peat deposits, palsas have degraded by 43% whereas lithalsas have decreased by 71%.

Because of the large extent of permafrost plateaus in the western part (shrub tundra) of the basin, it is impossible to count the exact number of mounds and assess how many individual landforms have disappeared. In the forest tundra, the number of palsas had decreased by 55%. The average area occupied by a lithalsa and a palsa was 1,806 and 954 m2 respectively in 1957. These values have decreased to 1,274 and 895 m2 in 2009 (Table 2.4). No new permafrost mound, either palsa or lithalsa, formed during this period.

Statistics on landslides and gullies are presented in Table 2.5. Of the 38 features identified in 2009, 12 were clearly present and active in 1957 and 2009, 12 of the landforms were clearly absent in 1957, and 14 of the possible already active features in 1957 were unidentifiable because of the presence of snow patches on the aerial photographs. The number of active landslides and gullies increased from 46 to 217% between 1957 and 2009 (Table 2.6). This large percentage range is due to the residual snow covers masking some slopes on the 1957 air photographs.

Location Biome Lenght Horizontal Height Breath Slope Surface Depth Volume Mass Present distance in 1957?

m m m m ° m2 m m3 t Landslides 76°27'55'' W / 56°37'28'' N ST a 30 7 8 11 50 327 - 876 911 no 76°20'50'' W / 56°37' N ST 32 5 6 8 50 249 - 477 549 yes 76°20'47'' W / 56°37'1'' N ST 35 10 5 11 27 393 - 892 1137 yes 76°19'44'' W / 56°36'56'' N ST 24 4 4 6 45 136 - 192 258 no 76°20' W / 56°36'57'' N ST 21 4 7 8 60 168 - 291 312 yes 76°20'42'' W / 56°37'1'' N ST 14 5 6 8 50 109 - 209 240 no 76°19'4'' W / 56°37'15'' N FT b 22 18 21 28 50 616 - 4247 3581 yes 76°19'7'' W / 56°37'13'' N FT 68 15 18 23 50 1587 - 9117 7931 yes 76°19'21'' W / 56°37'7'' N FT 26 3 5 6 60 156 - 203 235 no 76°19'20'' W / 56°37'6'' N FT 27 4 7 8 60 216 - 374 401 no Active layer detachment failures 76°26'38'' W / 56°37'22'' N ST 19 18 15 23 40 567 0.75 425 746 yes 76°26'36'' W / 56°37'25'' N ST 32 24 20 31 40 1152 0.75 864 1516 yes 76°19'50'' W / 56°36'56'' N ST 15 27 23 35 40 529 0.75 397 696 no 76°23'5'' W / 56°37'16'' N ST 10 18 15 23 40 235 0.75 176 309 snow 76°24'48'' W / 56°37'28'' N ST 28 17 14 22 40 621 0.75 466 818 yes 76°19'46'' W / 56°37'1'' N ST 23 10 8 13 40 300 0.75 225 395 snow 76°23'16'' W / 56°37'18'' N ST 11 25 21 33 40 359 0.75 269 473 no 76°16'38'' W / 56°37'21'' N FT 12 12 10 16 40 188 0.75 141 247 no Gullies 76°24'2'' W / 56°37'19'' N ST 18 38 32 50 40 893 0.9 804 1410 yes 76°26'19'' W / 56°37'16'' N ST 156 28 28 40 45 6177 0.9 5560 9757 yes 76°26'3'' W / 56°37'21'' N ST 25 10 10 14 45 354 0.9 318 558 no 76°25'58'' W / 56°37'15'' N ST 20 11 11 16 45 311 0.9 280 491 no 76°25'52'' W / 56°37'14'' N ST 38 21 21 30 45 1129 0.9 1016 1783 yes 76°25'15'' W / 56°37'17'' N ST 40 10 10 14 45 566 0.9 509 893 snow 76°20'38'' W / 56°37'4'' N ST 21 12 12 17 45 356 0.9 321 563 snow 76°20'4'' W / 56°36'52'' N ST 95 14 14 20 45 1881 0.9 1693 2971 yes 76°20'22'' W / 56°37'37'' N ST 82 4 4 6 45 464 0.9 417 733 snow 76°20'9'' W / 56°37'35'' N ST 18 6 6 8 45 153 0.9 137 241 no 76°19'55'' W / 56°37'36'' N ST 36 3 3 4 45 153 0.9 137 241 snow 76°20'33'' W / 56°37'42'' N ST 51 8 8 11 45 577 0.9 519 911 snow 76°20'34'' W / 56°37'15'' N ST 65 8 8 11 45 735 0.9 662 1162 snow 76°23'36'' W / 56°36'30'' N ST 63 9 9 13 45 802 0.9 722 1267 snow 76°23'55'' W / 56°36'29'' N ST 60 11 11 16 45 933 0.9 840 1474 snow 76°23'36'' W / 56°36'25'' N ST 30 10 10 14 45 424 0.9 382 670 no 76°23'25'' W / 56°36'26'' N ST 42 7 7 10 45 416 0.9 374 657 snow 76°16'5'' W / 56°37'56'' N FT 67 5 5 7 45 474 0.9 426 748 snow 76°17'37'' W / 56°37'22'' N FT 20 11 9 14 40 287 0.9 258 454 snow 76°16'39'' W / 56°37'18'' N FT 38 9 9 13 45 484 0.9 435 764 snow Total 35,652 48,505

Table 2.5 Characteristics and measurements of the landslides, the active layer detachment failures and the gullies of the Sheldrake River catchment in 2009. a shrub tundra;b forest tundra

Number of Total eroded Eroded organic Characteristics in 1957 and in 2009 landforms sediment matter a m3 t m3 t

Absent in 1957, present in 2009 12 3778 5422 49 70 Present in 2009, snow covered in 1957 14 6023 10,571 78 137 Active in 1957 and 2009 12 25,851 32,512 336 423 Total in 2009 38 35,652 48,505 463 631 Variations 46

Table 2.6 Number, volume and mass changes in landslides and gullies between 1957 and 2009. a The mass and volume of eroded organic matter are included in the mass and volume of total eroded sediment.

2.7.2 Activity of landslides, active layer failures and gullies between 1957 and 2009

The activity of landslides, active layer failures and gullies increased by a factor of 1.12 to 1.38 between 1957 and 2009 (Table 2.6). A total of 35,652 m3of eroded sediment volume from active landslides and gullies was calculated in 2009, which corresponds to 48,505 t of fine sediments, with an average content of 1.3% of OM (Calmels and Allard, 2008), i.e. 463 m3 (631 t). The 12 newly active landslides (observed in 2009 only) correspond to an eroded volume of 3,778 m3 (5,422 t) of clay, of which 49 m3 (70 t) is of OM. The 14 gullies and active layer failures unidentifiable on the 1957 photographs correspond to an erosion of 6,023 m3 (10,571 t) of clay, which contains 78 m3 (137 t) of OM. Landforms already active in 1957 show a lost volume in 2009 of 25,851 m3 (32,512 t) of clay with 336 m3 (423 t) of OM. By averaging landslide activity over the study period, 686 m3.yr-1 (933 t.yr-1) of sediment and 9 m3.yr-1 (12 t.yr-1) of OM would have been released annually by permafrost- related landslides.

2.7.3 Hydrological connectivity

The drainage (i.e. stream and channel) density increased in the selected sub-basin from 13.2 in 1957 to 15.6 in 2009 (+18%) (Figure 2.8).

In 2009, 0.24 km2 of the 1.09 km2 covered by thermokarst ponds were connected to the main stream. 0.09 km2 of the new 0.54 km2 of water coverage, resulting from the pond

41 expansion between 1957 and 2009, were connected to the drainage network (NCTP) (Table 2.3).

Figure 2.8 Change in drainage density caused by lithalsas and palsas decay in land system 24.

2.7.4 Potential TSS, DOC and OM flow in the system from CTPs

From mid-June to mid-October, 21 rain events that increased water levels occurred in 2009; 22 in 2010 (Figure 2.7). The cumulative rainfall associated with these events is 279 mm in 2009 and 402 mm in 2010. Assuming that CTP coverage stayed roughly the same in 2009 and in 2010, i.e. 0.24 km2 for CTP and 0.09 km2 for NCTP, transfers through the network of thermokarst ponds and connectors into the fluvial system are estimated at 1.01 t in 2009 and 1.46 t in 2010 for TSS, and 0.34 t (2009) and 0.49 t (2010) for DOC. Out of these

amounts, 0.37 t in 2009 and 0.54 t in 2010 for TSS and 0.12 t and 0.18 t for DOC come from recent thermokarst ponds formed between 1957 and 2009 (Table 2.7).

Surface area Cumulated rain Water yield a TSS b yield DOC c yield km2 m m3 t t 2009

CTPd 0.242 0.279 67,380 1.01 0.34 NCTPe 0.089 0.279 24,795 0.37 0.12 2010

CTP 0.242 0.402 97,085 1.46 0.49 NCTP 0.089 0.402 35,726 0.54 0.18

Table 2.7 Yields of sediment in suspension and dissolved organic carbon in tons from connected thermokarst ponds to the drainage system, in 2009 and 2010. a the volume of water which overflows from the thermokarst ponds and joins the fluvial system is equal to the surface area of CTPd or NCTPe multiplied by the height of the cumulated rain falling between mid-June to mid-October 2009 and 2010; b Total suspended sediment; c Dissolved organic carbon; d Connected thermokarst pond; e Connected thermokarst pond created between 1957 and 2009.

In 2009, CTP resulting from the collapse of palsas covered 21,510 m2 more than in 1957 (Table 2.3). Assuming that this area had an average peat thickness of 1 m with a dry bulk density of 0.1 t/m3 (Robinson and Moore, 1999; Vitt et al., 2000b), a total of 21,510 m3 equivalent to 2,151 t of peat has been potentially mobilized into the drainage network as OM, i.e. POM and DOM, during the 52 years period. 96% of that amount was released from the forest tundra area. Assuming a constant rate, 41 t/year of OM have been released..

2.8 Discussion

2.8.1 Toward a permafrost-free landscape

In the Sheldrake River catchment, area covered by lithalsas and palsas has decayed by 21% and area covered by thermokarst ponds has increased by 96% in 52 years. But these average values conceal important differences within the catchment area. Permafrost has degraded roughly at the same rate (23%) as was observed for riparian palsas 150 km to the north near the tree line by Vallée and Payette (2007), but sensibly less than a palsa peatland

43 to the south, near the Lac Guillaume Deslile, where palsas decreased by 83% between 1957 and 2003 (Payette et al., 2004) or lithalsas 8 km south of the study area, i.e. 40% (Fortier and Aubé-Maurice, 2008) (Table 2.8).

Change in Change in permafrost thermokarst Studies Location Area Type of permafrost Period extent pond extent Payette et al., 2004 56°11'N/75°55W 0.15 km2 palsa peatland 1957-2004 -83% 48% b Fortier et Aubé-Maurice, 2008 56°33'N/76°30'Wa 2.25 km2 lithalsas 1957-2005 -40% 175% Marchildon, 2007 56°51'N/76°15W 3×1.5 km2 lithalsas and palsas 1957-2005 -36% 79% Vallée and Payette, 2007 57°45N/76°20'N 14 landforms lithalsas and palsas 1957-2001 -23% 76% This study The Sheldrake River watershed 56°37’N/76°32’W 76 km2 lithalsas and palsas 1957-2009 -21% 96% Forest tundra " 43 km2 lithalsas and palsas " -43% 103% Shrub tundra " 33 km2 lithalsas " -7% 73%

Table 2.8 Permafrost and thermokarst pond changes at different sites in northern Québec from references. a 9 sites around the village of Umiujaq; b Estimate from authors results.

The Hudson Bay is the principal factor driving the east-west climate gradient characterised by colder coastal temperatures than inland (Lévesque et al., 1988). The colder coastal zone is generally characterised by cooler summers due to frequent marine fog and high cloudiness and by cold winters due to strong and dry winds once the bay is frozen. The inland climate, which corresponds to the forest tundra at ~15 km from the coast (Figure 2.9), is warmer and characterised by a thicker and less dense snow cover, particularly in forest stands. As a result, permafrost is patchier and decays more rapidly inland, east of the tree line, in the forest tundra than in the coastal shrub tundra.

Several environmental factors such as vegetation and surficial geology are responsible for the observed differences in thawing rates and degradation features within the Sheldrake River catchment. On this 76 km2 watershed, the range of degradation varies from 1 to 100% depending on distance of Hudson Bay, and on topographic, hydrologic, ecological and geomorphic factors. Palsas have decayed less rapidly than lithalsas, likely because thermal properties of peat help slow down the permafrost thawing (Vitt et al., 1994).

Pond coverage expansion follows permafrost decay. However the rate of pond formation is less than the rate of permafrost decay. Whereas 1.14 km2 of permafrost has disappeared since 1957, only 0.54 km2 of thermokarst ponds has formed over the same period. Explanations are: first, a pond replacing a lithalsa has a smaller diameter than the original mound due to the preservation of the rim ridge. For example, a 196 m2 typical lithalsa (50 m in diameter) was replaced by a 152 m2 pond with a 44 m2 rampart (3 m in width). Second, terrestrialization of thermokarst ponds due to colonization by sedges and mosses can reduce the rate of thermokarst pond coverage (Zuidhoff and Kolstrup, 2000; Payette et al., 2004). This is likely the reason why the area covered by ponds in peatlands is

45 proportionally less extensive than in the rest of the catchment. Third, permafrost fragmentation, talik formation and subsequent drainage can lead to shrinking lakes and disappearance of thermokarst ponds through rapid (even catastrophic) drainage (Yoshikawa and Hinzman, 2003; Smith et al., 2005; Jones et al., 2011; Sannel and Kuhry, 2011). Finally, in well drained terrain, evidence of permafrost disappearance can be associated with vegetation growth (shrub) in hollows, while in wet areas or when a lithalsa or a palsa collapses on the banks of the river, a stream or a lake, it merges with it and is not considered anymore as a thermokarst feature. On the other hand, ponds can also merge to form larger thermokarst lakes (Seguin and Allard, 1984).

As mentioned by Vallée and Payette (2007), the area covered by thermokarst ponds increased more rapidly than the number of new ponds (96% versus 50%), indicating that a number of ponds in 1957 were associated with non-totally thawed permafrost mounds or plateaus. Since then, the lithalsa or palsa has totally disappeared and the pond occupies a larger area. However, the increasing number of lithalsa landforms and their decreasing individual average size in the forest tundra area indicate a general fragmentation of the permafrost.

In 1957, thermokarst ponds covered 2.6% of surficial deposits for a cumulated area of almost 0.56 km2. Permafrost was then slowly degrading (Allard and Seguin, 1987b). The 1.4°C warming between the end of the LIA and the 1940’s in Northern Québec (Chouinard et al., 2007) is likely responsible for this early degradation. Permafrost probably reached its maximum expansion during the LIA and then started to decay (Allard et al., 1987; Payette et al., 2004; Arlen-Pouillot and Bhiry, 2005). After a short and light cooling between 1940 and 1990, while permafrost continued to thaw essentially due to a rise in snow precipitation (Payette et al., 2004), air temperatures have increased suddenly by 1.7°C until 2005 (Chouinard et al., 2007). In 2010, an average annual air temperature of 0.0°C was recorded at Umiujaq, whereas the long term average of the annual air temperature noted by Environment Canada (1925-2010) is -5°C. This supports observations of a recent acceleration of the permafrost degradation processes in the forest tundra biome (Payette et al., 2004). In areas where the losses exceeded 50% between 1957 and 2009, permafrost would have completely disappeared if the rate of degradation would have been continuous

since the end of the LIA. Land system 31 is now devoid of permafrost. Considering that the increase of precipitation is predicted to persist and that temperature will continue to warm up by about 3 to 4°C until 2050 (Allard et al., 2007; IPCC, 2007; Brown et al., 2012), it is anticipated that most of the permafrost shall disappear in the forest tundra in the coming decades. Only fossil thermokarst features will remain.

2.8.2 Importance of the activity of landslides, active layer failures and gullies

Landslides are induced by fluvial erosion, which creates active layer slumping in the ice- rich permafrost and accelerates thaw and thickening of the exposed active layer. In summer, clay liquefaction leads to failure of the “newly exposed active layer”. The headwall recedes by slumping and remains parallel to the original slope angle. However, these landslides do not expose the underlying permafrost since the active layer on the slope is too deep. The landslides are concentrated along steep river banks, particularly in areas of permafrost plateaus, in the western part of the studied basin (see photograph on Figure 2.6). The released sediments are rapidly evacuated by river flow. No ground ice exposure was ever observed despite the fact that drilling only a few meters behind the cliff head revealed abundant ground ice (Allard et al., 1996).

Most of these landslides show signs of high activity. In summer 2009 and 2010, the four landslides showed blocks of turf and clay that had recently slid from the top to the bottom of the slope. The smooth and water saturated surface of the headwall and the absence of vegetation cover also indicate that erosion is active. A new landslide occured during the thaw period (June) of 2010 (Figure 2.4f). The three landslides having the gentlest slopes (25-27°) were already active in 1957. Three slides now appear to be entering in their stabilisation phase as their slope is becoming too gentle to trigger detachment and the permafrost table has retreated several meters from the top of the headwall which is rapidly colonized by shrubs.

Active layer failures occur on the steep (35-40°) upper edge of permafrost plateaus (Figure 2.4d). They are scattered throughout the study area. In 1957, three of them were already active, three were absent and two were covered by snow. The two failures visited in 2009 show limited activity. As an active layer failure is triggered by annual thawing and since

47 the slide is confined into the top part of the active layer, as noticed by Seguin and Allard (1984), the erosion form is maintained active by two processes. First, melting of snow at the end of spring and rain in summer over-saturate the surface layer of the soil and lead to liquefaction of clay, which flows on the slope. Second, frostboils oriented across the slope continuously bring liquefied or viscous fine sediments to the surface and can slide and trigger small active layer failures (Figure 2.4b).

Gullies are widespread landforms in the watershed, either singly or as a network. Generally, gullies have symmetric slope angles along the edges of two elongated plateaus. They deepen and widen during rain events. The slope of the permafrost plateau edge increases (40-45°) leading to skin failures, which grow and deepen year after year. 80 % of the listed gullies were already present or covered by snow in 1957. Several former gullies now (and in 1957) colonized by a thin layer of vegetation show that incision between plateaus is not a recent process. As for active layer failures, it is likely that these erosional landforms were present in the basin since the LIA. It is also possible that some plateaus have been divided in several smaller individual mounds by gullying. Nevertheless, the gullies are still very active and should continue to erode and export significant volume of sediments during future rain events. Almost two gullies out of three were covered by snow in 1957. This suggests that snow accumulation between plateaus reduces frost penetration in winter and contributes to trigger gully formation, as emphasised in several studies on palsa collapse (Seppälä, 1990, 1994; Matthews et al., 1997) and degradation of permafrost plateaus (Allard et al., 1996).

2.8.3 Increase of sediments and OM released by landslides, active layer failures and gullies

The higher number of landslides observed in 2009 (+46 to +217%) suggests that the overall geomorphic activity on slopes and in gullies has increased, resulting in higher volumes of eroded sediments in 2009 than in 1957 (+12 to +38%). It is in accordance with several other studies, which noticed an increasing number of active thaw slumps in the Mackenzie River delta (Lantz and Kokelj, 2008), in Hershel Island (Lantuit and Pollard, 2008) and on the Aklavik Plateau (Lacelle et al., 2010). However, it seems that the small size of

landslides and gullies results in small inputs in the Sheldrake River fluvial system. The total volume of eroded sediments in the catchment from the 38 erosion landforms related to permafrost disturbances is only 35,652 m3. For comparison, changes in gully size in the only non permafrost complex of erosion in the center of the basin (land system 16) suggest a loss of 26,600 m3 of clay. The only study in Northern Québec that quantifies landslide scars reveals that 7 rotational landslides that occurred in unfrozen clay of the Great Whale River estuary have exported 17×106 m3 of sediments (Begin and Fillion, 1987). Thaw slumps in continuous permafrost can affect several thousands of m2 of terrain (Lantuit and Pollard, 2008). Therefore, landslides induced by permafrost decay in the Sheldrake River basin are a small source of sediment in comparison with erosion landforms and mass wasting assessed elsewhere, both in permafrost and in non permafrost areas.

The volume of lost permafrost terrain in about 50 years inevitably means that sediment and carbon amounts released in the fluvial ecosystem have increased. It is likely that a part of these products is temporarily stored in the aquatic system, as bottom sediments in ponds and streams. However, eroded sediment and carbon will ultimately reach Hudson Bay.

2.8.4 Impact of an increasing connectivity on organic carbon and mineral sediment yields

The important increase in catchment internal connectivity, reflected by increases of 58% in CTP and 18% in stream density (Figure 2.8) from 1957 to 2009, represents a major environmental and hydrological change. Despite large methodological uncertainties, degradation of permafrost and the development of thermokarst ponds currently release a significantly higher amount of DOC and TSS in the fluvial ecosystem than previously. No comparison with other studies is available yet to our knowledge.

This increasing connectivity also results in a more efficient drainage network and a fragmentation of the remaining permafrost. Consequently, in discontinuous permafrost, as in the continuous permafrost (Rowland, 2010), recent climate warming has altered the water balance. Thermokarst was already occurring in 1957. Results clearly indicate that the rate of permafrost decay has recently increased in the studied area, probably since 1990, as observed elsewhere in the region (Payette et al., 2004; Fortier and Aubé-Maurice, 2008).

This increasing rate of permafrost degradation probably occurred following the significant increase of air temperature in north-eastern Canada during the 1990’s (Chouinard et al., 2007).

CTPs in the forest tundra area export 2.3 times more TSS and DOC per unit of area than the CTPs in the shrub tundra area. This amount increases to a factor of 2.6 for NCTPs. The first explanation is likely the difference in the ratio of the CTP coverage per km2 that is 3.6 times superior in the forest tundra. Conversely, landslides, active layer failures and gullies are more abundant in the shrub tundra area. They currently contribute 70% of the total sediment and carbon released by landslides in the river catchment. Therefore, the tree line is a major boundary between two regimes of thermokarst processes.

The measurement of CTP inputs based on rain events was done in summer conditions only. Thus, this approach does not include snowmelt and break up during which, ponds overflow and hydrological connectivity is temporarily higher (Bowling, 2003; Woo and Guan, 2006). An important fraction of annual POM and DOC is likely transferred through the fluvial system during the melting period. For example, in the Great Whale River, 100 km south of the area, 23% of annual POC and 11% of annual DOC are exported during the break up (Hudon et al., 1996).

The primary source of old POM in the basin is the peat released from the decay of palsas which have degraded by 55% since 1957 (Tarnocai et al, 2009; Kuhry et al., 2010). Not all of the released OM from palsa decay may reach the fluvial network. However, it is likely that a non-negligible fraction of the eroded peat, which collapsed in CTP, is temporarily stored in the ponds to be thereafter exported in the river by overflow during rain events.

2.9 Conclusion

Permafrost has considerably degraded east of Hudson Bay, as elsewhere in the discontinuous zone in circum-arctic regions. Palsas and lithalsas of the Sheldrake River catchment have generally decayed by 21% between 1957 and 2009. The rate of thaw was much higher in the catchment head (forest tundra area) than near the basin outlet on Hudson Bay coast (shrub tundra area). The tree line is a definite climate and ecological

boundary that significantly impacts on the rate of permafrost decay and on the volume of sediment and carbon yields. Climatic projections for the next decades indicate that this trend will continue until permafrost disappears.

Increase in activity of landslides and in thermokarst pond cover between 1957 and 2009 resulted in significant sediment and carbon release in the fluvial system. However, the total computed amounts of currently released sediments and carbon from the Sheldrake River basin remains small compared to the Great Whale River, a permafrost free catchment where landslides occurred in the same type of clay deposits. This illustrates that beyond a certain stage of degradation (now, 20% left above the tree line and 10% below), the impact of residual permafrost degradation on sediments and carbon release in fluvial systems declines quickly.

2.10 Acknowledgements

This project was supported by grants from Arcticnet and the Natural Science and Engineering Research Council of Canada (discovery grant to M. Allard). The Centre d’études nordiques of Université Laval provided important logistical support. We thank Denis Sarrazin for the installation and maintenance of all the data gathering equipment in the study region. The field assistance of Catherine Falardeau-Marcoux and the help in mapping by René-Charles Bernier, Emmanuel L’Hérault and Carl Barrette were appreciated. The suggestions of Warwick Vincent and Mickael Lemay on a preliminary version of the manuscript are appreciatively acknowledged. The paper was significantly improved by the reviewing process. We are also grateful to the Inuit community of Umiujaq for its generous hospitality.

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CHAPITRE 3 - Hydrological regime and sediment transport in a river flowing in a thermokarst landscape, Sheldrake River, Nunavik, Quebec.

3.1 Résumé

Le régime hydrologique et le transport de sédiment dans les rivières subarctiques restent peu documentés, principalement parce que la difficulté d’accès dans ces régions requière une logistique et une instrumentation complexes. Une station de jaugeage, installée en 2008 sur la rivière Sheldrake, dans la zone de pergélisol discontinu de la côte est de la baie d’Hudson, a fourni des données de niveau d’eau et de turbidité. Des prélèvements d’eau et des observations de terrain ont également été réalisés. Le débit fluvial et les concentrations en sédiments furent mesurés. Le débit de la rivière Sheldrake, typique des régimes hydrologiques subarctiques, est principalement influencé par la fonte des neiges et régulé par les précipitations estivales. La turbidité varie avec les fluctuations de débit liées aux précipitations estivales. Cependant, l’étude montre que la charge sédimentaire maximale est indépendante de la crue printanière. En 2010 et 2013, respectivement 53 et 25% des sédiments transportés entre avril et novembre le furent pendant la période de bas niveaux d’eau, en juillet, alors que les températures de l’air estivales provoquaient une rapide pénétration du front de dégel dans la couche active du pergélisol, atteignant 85 à 90% de sa profondeur maximale. Par conséquent, il est suggéré que le dégel de la partie basale de la couche active, riche en glace, a engendré une activation des ostioles et favorisé le déclenchement de glissements de terrain, à l’échelle du bassin versant. Ceci a entrainé une augmentation de l’érosion des sols et du volume de sédiments libéré dans le système fluvial. L’étude démontre ainsi que le débit solide maximal d’une rivière dans une région de pergélisol peut être contrôlé par la température de l’air et non strictement par le régime hydrologique.

3.2 Abstract

The hydrological regime and the transport of sediments in subarctic rivers still remain poorly documented, mostly because of their remoteness and the complex logistics required

59 to install monitoring instrumentation. A gauging station installed in 2008 on the Sheldrake River located in the discontinuous permafrost zone on the east coast of Hudson Bay, provided water level and turbidity data. Water sampling and observations of geomorphological processes in the river catchment were also done. River discharge and sediment concentrations were measured. Discharge follows a typical subarctic hydrological regime mainly driven by spring flood and regulated by rain events during summer. Turbidity varies with discharge variations regulated by precipitations during summer. The study reveals, however, that the maximum sediment load is not regulated by the river discharge regime. In 2010 and 2013, respectively 53 and 25% of the sediment transported between April and November was exported during low river stages in July when warm summer temperatures were causing the fastest rate of progression of the thaw front in the active layer of the permafrost, reaching about 85-90% of maximum thaw depth. It is therefore hypothesized that the thawing of the ice rich basal levels of the active layer leads to the activation of frostboils and triggers landslides throughout the river catchment, thus increasing soil erosion and raising sediment delivery into the fluvial system. The study indicates that maximum sediment discharge in a permafrost region may be predominantly thermally-driven rather than hydrologically-driven.

3.3 Introduction

The hydrological regime of subarctic rivers is strongly linked to seasonal climate variations that generate a large range in water discharge. Consequently, sediments loads are likely extremely variable. The hydrological cycle of subarctic/arctic rivers is regulated by snow storage and melting and freezing of soil water. Permafrost is a major factor that restricts infiltration and percolation at depth and maintains a perched water table near the surface in summer (Carey and Woo, 2001; Carey and Quinton, 2005; Quinton and Carey, 2008). Baseflow may cease in winter since sub-permafrost groundwater may be non-existent or too deep to discharge in the catchment, or because taliks can be only poorly connected with springs on the river beds. Permafrost warming, thinning and decay, earlier breakups and freshets, decline of the snow cover duration and increase in shrub cover are factors affecting the hydrology of high latitude rivers.

Thawing of permafrost potentially releases important volumes of sediments through thermokarst processes such as thaw slumping and thermal erosion. Sediments are mobilized in water courses and feed sedimentation in lakes and coastal seas (Guo et al., 2004; Goni et al., 2005). As thermokarst processes occur in a stochastic way in the landscape, releases can be expected to be variable in time and intensity.

Evidences of general permafrost decay have been observed in all high-latitude regions (e.g. Solid and Sorbel, 1998; Luoto and Seppälä, 2003; Jorgenson et al., 2006). In northern Québec, near the southern limit of permafrost distribution, thawing of permafrost in large areas of palsas, lithalsas, peat plateaus and permafrost plateaus has led to the disappearance of roughly 40% of permafrost extent over the last 50 years (Payette et al., 2004; Marchildon, 2007; Vallée and Payette, 2007; Fortier and Aubé-Maurice, 2008; Jolivel and Allard, 2013). Continuing warming is likely to lead to further degradation, making available previously frozen organic matter for bacterial decomposition and recycling into bio-available carbon and greenhouse gases (Schuur et al., 2008).

Deepening of the active layer, terrain subsidence, new and enlarging thermokarst ponds and landslides are the main evidences of permafrost degradation. These erosion features can lead to events of extreme sediment loading so far poorly reported in the literature (Lewis et al., 2005; Lamoureux and Lafrenière, 2009). A growing volume of sediments, organic carbon and other materials is potentially released by more numerous and active thermokarst features such as active layer slides and retrogressive thaw slumps (Kokelj et al., 2002; Lewkowicz and Harris, 2005; Jorgenson et al., 2006; Lantuit and Pollard, 2008; Lantz and Kokelj, 2008; Lacelle et al., 2010; Jolivel and Allard, 2013). These inputs can alter terrestrial and aquatic ecosystems and affect food webs as well as primary and secondary production (Kokelj et al., 2002, 2009; Bowden et al., 2008; Mesquita et al., 2010).

In Alaska, Toniolo et al. (2009) have detailed sediment transport and geomorphological processes in the spatio-temporal evolution of a thermokarst terrain. Other papers describe hydrological and sediment transport processes of watersheds underlain at different extents by permafrost (Kane et al., 2000; Hinzman et al., 2003; Toniolo and Kondial, 2006). However, measurements on rivers as vectors of geomorphic changes of the watershed in a context of permafrost decay still have received little consideration (Bowden et al., 2008).

Sediment transport and hydrologic response in a decaying landscape of palsas and lithalsas have not been studied yet, likely because those landscapes are remote and monitoring instrumentation has to be installed in harsh environments and in rivers with extremely variable discharge.

In this study, an analysis of hydrological and sedimentological regimes at the catchment scale is conducted in order to test for the hypothesis that the amount and timing of sediment fluxes in a subarctic stream are associated with the rate of thawing and the rate of geomorphic activity of the permafrost and thermokarst processes. The main objectives are to (1) document the hydrologic functioning and response to climate variations of a subarctic river in discontinuous permafrost, (2) provide an estimated budget of sediment in this river that flows in a thermokarst affected catchment, (3) describe two events of high turbidity related to periglacial and thermokarst processes, and finally (4) estimate the part of fluvial inputs originating from thermokarst in the total fluvial budget.

3.4 Study area

Located 30 km south of the Nastapoka River, on the eastern coast of Hudson Bay, the 25 km long Sheldrake River drains a 76 km2 watershed (Figure 3.1A). The river originates from Sheldrake Lake, on the Archean sector of the Canadian Shield. Near the coast, the river valley cuts through a range of coastal hills in Late Proterozoic bedrock and flows into Hudson Bay at 56°37’N; 76°32’W. On a topographic 1: 50,000 map, the low-gradient Sheldrake River is a second order stream. However, the rapid and recent permafrost decay makes the network of water tracks more elaborate and is currently increasing stream density (Jolivel and Allard, 2013).

The east-west elongated shape of the catchment (gravelius index: 1.9) is primarily due to the erosive activity of Pleistocene glaciers. Inland, the topography is dominated by flat valley floors, scattered with lakes and small hills with a general elevation of 200 to 250 m a.s.l.. The mean longitudinal river slope is 0.6%. It is <0.5% inland and increases to 3% once crossing the coastal hills towards the Hudson Bay.

The entire basin was submerged by the postglacial Tyrrell Sea from 8200 BP. Therefore, 85% of the surficial deposits of the area are marine silty clays (Figure 3.1B). Inland, sand and gravel deposits are associated with ice-contact glacio-fluvial deltas which mark the postglacial marine limit, whereas sand and gravel anchored on the slopes of the coastal hills have been reworked into beaches and terraces during the fall of the relative sea level (Figure 3.1B) (Lavoie et al., 2012; Lajeunesse and Allard, 2003). Because of permafrost, dispersed vegetation cover, presence of impermeable clay and bedrock, water infiltration and percolation rates are very low and most of water input from rainfalls and snowmelt flow as surface run-off. However, numerous wetlands, topographic depressions, thermokarst ponds and lakes characterizing the Sheldrake River catchment act as surface water storage areas.

The climate is subarctic, characterised by cold winters (-24°C in January), cool summers (10°C in August) and a mean annual temperatures varying between -4°C and -5°C. The area is covered by snow for ~ 8 months every year. Water inputs come from rain (60%) and snow (40%). The average annual evapotranspiration is ~ 200 mm.yr-1 (Payette and Rochefort, 2001). The eastern sector of Hudson Bay is generally ice-covered from early December to the end of May or beginning of June. However, during the winter 2010-2011, freeze-up did not occur until mid-January. The Sheldrake River has a ~1 m thick ice cover from early November onwards. The breakup occurs a few weeks before the melt of the Hudson Bay ice cover. Snowmelt generally occurs in late May and June but some snow banks can last until mid-summer. Between 2009 and 2013, the average date of the beginning of the Sheldrake River ice breakup was early May and the level of the river remained high during several weeks before drawing down to its summer flow (Table 3.1).

The basin is located at the transition between the sporadic permafrost zone and the widespread discontinuous permafrost zone in an area of high concentration of lithalsas, palsas and permafrost plateaus (Allard and Seguin, 1987). Permafrost is present in 20% of the surficial deposits (Jolivel and Allard, 2013). The tree line crosses the basin and is oriented from south to north due to the cooling influence of the Hudson Bay; the landscape close to the bay is dominated by the shrub tundra while the forest tundra extends further inland (Figure 3.1B). In the western part of the basin (shrub tundra area), the river and its

tributaries flow through an area of permafrost plateaus, incising meanders and deep tributary gullies. In the forest tundra, some peat deposits over clay or sand extend over flat and poorly drained valley bottoms. Most of the palsas in the catchment are found in these bogs.

The Sheldrake river catchment is an area of intensive permafrost decay. Between 1957 and 2009, 21% of the permafrost has thawed. This degradation is more important inland where some areas are now devoid of permafrost. As a result, numerous thermokarst features, such as thermokarst ponds, landslides, active layer failures and gullies, release a growing volume of sediment into the Sheldrake river system. Thus, between 1957 and 2009, the increase of thermokarst pond cover and of stream connectivity related with permafrost fragmentation allowed for an increase of sediments fluxes to the main stream by a factor of 1.6. The activity of landslides and gullies also increased by 12 to 38% (Jolivel and Allard, 2013).

3.5 Methods

3.5.1 Field instrumentation and laboratory analyses

A gauging station was installed 2 km upstream from the river mouth in late summer 2008 (see Figure 3.1B for location). As the river bed is rocky and the current is swift in the lower reach of the river, the instrumentation had to be installed between two rapids in a convenient pool of calm water. This automated station continually records water temperature, stage and turbidity. The mooring consists of a dead weight (20 kg) attached to a buoy that rests on the river bed and is tied to the shoreline with a steel cable. The buoy was submerged under ~ 1.2 m water at a distance of 4 m from the river bank in order to avoid being swept away during ice breakup. The data is retrieved three times a year (generally in June, August and October) due to the remoteness of the site. The instrumentation attached to the mooring consists of a Levellogger (Solinst) and an OBS 3+ (Optical Backscatter Sensor, Campbell). The levelogger was calibrated for barometric compensation with a barologger (Solinst). The mooring is situated just upstream a 20 m waterfall. Thus, at spring, the breakup is facilitated even if the upper part of the river break up several days later.

Discharge records were generated by rating the river using the velocity-area method. Stream velocity was measured directly on the site from a small inflatable boat with a current velocity meter 2100-LX. Twenty sets of measurements were taken at two meters intervals along a transverse section in front of the station and at three depths during different flow levels. Velocity measurements were made in summer 2010, 2011 and 2013 and combined to produce a stage/discharge curve.

The OBS 3+ uses its sideways-facing optics to emit a near-infrared light into the water. It then measures the light that bounces back from the water’s suspended particles. The OBS 3+ sensor head was cleaned at every site visit while servicing the station to prevent accumulation of algae that could have invalidated the data (Schoellhamer, 1993). Analyses were made by the Laboratoire de l’INRS, Centre Eau, Terre, Envionnement and by the Environex laboratory of Quebec City. TSS was measured from filtration through a 0.45 μm filter and weighting after drying.

Air temperature and precipitation data were provided by an automated meteorological station, operated by the Centre d’études nordiques, located 8 km south of the Sheldrake River, near the village of Umiujaq.

3.5.2 Stage/discharge calibration

Since the transverse section is rather regular, a logarithmic curve was applied to the data for determining the stage-discharge relationship (Figure 3.2a):

Q = 0.058e 0.021 Level (1)

With r2 = 0.9 P < 0.001 S= 2.6 N= 20

Even if the correlation is statistically significant, the small number of velocity measurements makes results less accurate for extreme flows. The calibration is valid only during the ice free period. When there is an ice cover, pressure and roughness of the ice and snow cover, make the levellogger inaccurate. In this study, according to hydrographs (Figure 3.3), the freshet flood event is considered as the period during snowmelt period (several consecutive days) when the discharge exceeds 20 m3/s. The end of snowmelt time

is associated with the resumption of river flow to its summer stage (i.e. at 12 m3/s), and when air temperature is maintained above 0°C.

3.5.3 NTU/TSS calibration

The NTU/TSS relationship is given by the following equation:

TSS = 0.46NTU + 2.457 N=22 (2)

With r2= 0.88 P < 0.001 S= 1.13

Due to remoteness of the site, only three samples in the calibration data illustrate conditions of relatively high TSS (total suspended sediment) concentration (≥14 mg/L) (Figure 3.2b). Despite this fact, the coefficient of determination (r2) and the probability value (P) of the linear regression are significant and allow discussion about sediment transit. Changes in sediment transport in this study are expressed in NTU (nephelometric turbidity unit) variations. The NTU/TSS curve is used only to provide a rough estimate of sediment transport. Further, the cable linking the submerged instrument and the datalogger on shore was severed by ice and complete summer coverage of NTU was obtained only in 2010 and 2013 (until 19 October). As for discharge, turbidity data are expressed as mean daily values.

The relationship between NTU and TSS concentration is almost linear (Downing, 2006), particularly for fine sediment (Lewis, 1996), and can provide continuous data of sediment transport in remote access rivers, such as the Sheldrake.

Figure 3.2 (a) Relationship between water level (m) and discharge (m3/s); (b) relationship between turbidity (NTU) and total suspended sediment (TSS) (g/m3).

3.5.4 Active layer depth

The evolution of the active layer depth in 2010 was estimated thanks to thermal data coming from conventional thermistors attached on a cable and installed in a borehole drilled in permafrost. Data were download every summer since 2002 and run with a matlab program. Unfortunately, no data are available for 2013.

3.6 Results

3.6.1 Interannual and seasonal variations of discharge

Figure 3.3 shows the recorded hydrographs from 2009 to 2013; corresponding hydrological features are detailed in Table 3.1. Average daily discharge during the ice free period varies

from 13 to 22 m3/s. The freshet high stages last between 16 and 45 days. The 2012 flood is the more abrupt one of the five years with an extreme peak discharge of 142 m3/s reached on 7 June. In 2012, 44% of the ice free period discharge flowed during the spring flood, while this value was around 20% in 2010 and 2011.

The snowmelt period lasted for 7 to 9 weeks, except in 2010 (11 weeks) which had the earliest snowmelt start date of the five years of the study period. The Sheldrake River reaches its summer flow at the end of June or at the beginning of July. In autumn, successive rainfall events generate an increase in discharge from September until the river freezes (Figure 3.3).

160 2009 2010 2011 2012 2013 140 120 100 80 60 40 20 0 23-Apr 23-May 23-Jun 23-Jul 23-Aug 23-Sep 23-Oct 23-Nov

Figure 3.3 Hydrographs for 2009, 2010, 2011, 2012 and 2013. Curve of discharge starts when discharge reaches 12 m3/s, which is supposed to be the first signs of ice break up.

Average Length Length Peak % freshet flood in Date start/end Start/end discharge snowmelt breakup discharge summer snowmelt freshet flood (m3/s) (days) (days) (m3/s) hydrological balance

2009 17 15 May/9 July 56 7 June/1 July 25 69 (14 June) 34 11 May/1 2010 16 24 April/18 June 76 22 46 (27 May) 20 June 4 June/19 2011 13 11 May/26 June 47 16 46 (5 June) 19 June 24 May/29 2012 22 3 May/7 june 66 37 142 (7 June) 44 June 1 May/14 2013 16 1 May/22 june 53 45 66 (3 May) 40 June

Table 3.1 Hydrological features of the Sheldrake River in 2009, 2010, 2011, 2012 and 2013 during the ice free period.

3.6.2 Turbidity regime over the thawing season

3.6.2.1 Turbidity variations during the snowmelt period

In 2010, the snowmelt period lasted from 24 April to 18 June. During this period, daily NTU values varied between 5 and 15 and peaked at 22 on 6 May (Figure 3.4). This value was reached after five days with air temperatures fluctuating ±0°C and some days before the true beginning of the freshet flood (11 May). During the spring flood, turbidity values did not increase significantly. However, in early June, five days after the maximum of the freshet flood, turbidity increased and stayed at a level of ~7 NTU, compared with a mean 4 NTU in May. This is probably due to the melting of remaining snow banks in the landscape and melting of snow and ice in lakes.

In 2013, the period of high water stages lasted 45 days and ended at mid-June (Figure 3.5). Daily NTU values generally varied between 3 and 20 NTU. Contrary to 2010, high discharges caused by breakup and snowmelt, are associated with a higher turbidity, ranging from 14 to 20 NTU from 1 May to 5 May. From 11 to 14 May, turbidity increased significantly until reaching 135 NTU the 12 May. No precipitation (the most recent precipitation occurred on 6 May), no change in air temperature, which stayed negative, and no rise in river flow are noticed. It is possible that this peak is an artefact caused by the break up or by an ice jam in the pool.

From 8 to 12 June 2013, turbidity averaged at 59 NTU. At this time, most of the snow has disappeared in the landscape and the Sheldrake River water level lowered to reach its summer level at the end of June. Before and during this period, no significant rain was recorded (4 mm the 10 June). However, daily average air temperature raised from 0°C the 6 June to 14°C the 8 June. It maintained between 15°C and 18°C until 18 June.

3.6.2.2 Relation between turbidity and summer air temperature (2010 and 2013)

In 2010, the snow melt high levels did not influence significantly the suspended sediment transport. Whereas the period of maximum discharge occurs during the spring freshet, suspended sediment transport in the Sheldrake River reaches its summer maximum in July. Between 30 June and 23 July 2010 (24 days), NTU values averaged at 83 with a peak at 160 on 10 July, while the rest of the 2010 thawing period registered an average turbidity of only 5 NTU (Figure 3.4, Table 3.2). Results clearly show that there was a lag between the freshet flood (May-June) and the highest turbidity period that occurred in July (Figure 3.4). Moreover, no precipitation event exceeding 8 mm of rain occurred during this turbidity

period. On the other hand, the NTU curve rises three days following the increase of the average daily air temperatures from 0-5 °C up to a minimum of 10 °C for the rest of the turbidity period.

In 2013, a similar increase in turbidity occurred in midsummer but with a lower magnitude (Figure 3.5). The period of high turbidity lasted 19 days, i.e. 15 July to 2 August, with daily average NTU of 33, while the rest of the 2013 thawing period registered an average turbidity of only 7 NTU (Table 3.2). During this period, 10 days with light to moderate rain were registered (4 to 13 mm/day). The two main peaks of turbidity, i.e. on 17 July and 29 July, are associated with respectively 2 and 3 rainy days (10 and 25 mm). As a consequence, discharge is also correlated with turbidity and form small peak events. Finally, air temperatures are variable but no significant trend clearly appeared before or during the midsummer turbidity period (Table 3.2).

Context fitting with the onset of the period of turbidity

Number of Daily turbidity Rain Discharge Air temperature Period of high turbidity days min./max./av.

30 June - 23 July 2010 24 24/160/43 No preceding rain event Decrease Increase from 0-5 to 13 C

⁰ 2 days of light rain (8 and 6 Generally rising above 15 July - 2 August 2013 19 15/57/33 mm) preceding the onset of Stable 0 C the turbidity period ⁰ Table 3.2 Comparison between the summer periods of turbidity in 2010 and in 2013.

Finally, from September 2010 and 2013, small peaks of turbidity are directly correlated with precipitation and discharge. Air temperatures drop until the freeze-up. No turbidity event like those recorded in spring and midsummer occurred, although water levels were higher (Figures 3.4 and 3.5).

3.6.3 Total suspended sediment export estimates

Mean concentrations of TSS are 8.2 g/m3 in 2010 and 7.8 g/m3 in 2013 (Table 3.3). In 2010 and 2013, the Sheldrake river exported respectively a calculated mass of 160 t and 105 t of suspended sediment, corresponding to a specific sediment load per unit of catchment area of 2.1 t.km2.yr-1 in 2010 and 1.5 t.km2.yr-1 in 2013 (Table 3.3). During the midsummer turbidity periods of 2010 and 2013, the Sheldrake River exported respectively 53 and 25% of its annual suspended sediment budget. Notice that in 2013, the study period ended the 19 October. The computation of TSS only takes into account the Sheldrake River catchment starting downstream from the outlet of the Sheldrake Lake (3.4 km2), which has a subcatchment of 91 km2. The lake is a sedimentation basin and it is difficult to assess its contribution in TSS because it has never been sampled. Therefore, TSS specific values must be considered as maximums.

Average Start of the Number of days Average Total TSS/km2 Cumulated Discharge ice free of the study TSS exported (t) rain (mm) (m3/s) period period (g/m3) TSS (t)

2010 16 3 April 226 8.2 160 2.1 440 2013 * 16 29 April 174 7.8 117 1.5 490

Table 3.3 Sediment exported by the Sheldrake River and specific values in 2010 and 2013. *In 2013, data acquisition stopped on 19 October.

3.7 Interpretation and discussion

3.7.1 Hydrological regime of a typical subarctic river

Hydrological patterns were similar over the five-years, but the magnitude and timing of discharge varied considerably. The flow regime of the Sheldrake River is typical of those of the subarctic environment and is closely linked to precipitation and air temperature. The discharge is high at the end of spring (snowmelt) and in autumn (heavy rain), contrasting with low flow in summer (July and August) and non-measurable flow in winter.

In winter, the frozen soils, the snow cover and the absence of runoff prevent erosion and discharge and particulate transport reach their lowest level (e.g. Hudon et al., 1996; Chikita et al., 2012). In spring, the rise in solar radiation and above freezing air temperature rapidly melts snow cover and results in a rapid increase in discharge (Figure 3.3, 3.4 and 3.5). Snowmelt runoff is the dominant hydrologic event (19 to 44% of summer discharge) (Table 3.1). River ice and most of the snow cover melt before the surface of the soil begins to thaw, minimizing infiltration and high surface runoff. On the Sheldrake river catchment, surface meltwater runoff, resumption of an unimpeded hydrological connectivity between thermokarst ponds and streams and increase of river flow velocity augment fluvial inputs (Jolivel and Allard, 2013).

In periglacial regions, runoff rates and discharge response time tend to wane over the summer period as snow depletes, seasonally frozen soils thaw and the active layer increases in thickness over permafrost areas, altering water infiltration, percolation and soil water storage (McManara, 1998; Kane et al., 2000; Carey and Woo, 2001; Hinzman et al., 2003; Woo et al., 2008). Our hydrographs and field observations suggest further that, after the freshet, summer river discharge depends on both storm intensity and water storage capacity of the basin in soils, ponds, lakes and wetlands, as observed in southern temperate Québec (James and Roulet, 2009).

3.7.2 Impact of active layer thawing and thermokarst processes on sediment annual regime

In 2010, an important lag between the freshet maximum discharge and the maximum turbidity in early summer was oberved. The same general pattern occurred in 2013: except from 8 to 12 June, turbidity values recorded from 15 July to 2 August also dominate the thawing season. In rivers flowing in permafrost-free watersheds (Hudon et al., 1996) or in arctic catchments (Kriet et al., 1992; Braun et al., 2000; Forbes and Lamoureux, 2005), transport is normally a function of discharge and dependent on summer rainfall and on the amount of snow accumulation during the previous winter.

The year 2010 was the warmest on record in Nunavik and globally, with an average annual air temperature up to 0,62°C above the average air temperature registered since 1880

(NOAA, 2011). In Umiujaq, 2010 mean annual air temperature was very close to the freezing point (-0,07°C) compared to the long term average of -5°C (Environment Canada, 2010) (Figure 3.6) .

The year 2013 was colder than 2010, with an annual air temperature of -3,5°C (Figure 3.6), particularly due to a severe winter. However, the starting of both 2010 and 2013 thawing seasons were the earliest in the study period, with first signs of snowmelt recorded respectively on 24 April and 1 May (Table 3.1).

1997 1999 2001 2003 2005 2007 2009 2011 2013 0

-1

-2

-3

-4

-5

-6

Figure 3.6 Annual air temperature in Umiujaq between 1998 and 2013. No data are available in 1999, 2001 and 2002.

In July 2010, the turbidity period coincides with the period of the summer when the thaw front reaches about 85-90% of the active layer depth (Figure 3.7). Based on the degree-days curve, the 2013 summer period of turbidity likely occurred at a similar timing in the thawing season than for 2010 (Figure 3.8). The basal part of the active layer in fine-grained soils such as clay is usually a zone of ice enrichment during winter freeze back (French, 2007). Rapid thawing of the basal ice rich layer can lead to slope detachment slides (Lamoureux and Lafrenière, 2009). The turbidity period also coincides with the steepest part of the cumulative thawing degree-days curve, i.e. when thaw front was still rapidly penetrating to the base of the active layer. This is the time of the year when frostboils on clay and silt rich soils are reactivated by excess pore water pressure caused by soil thawing, bringing to the surface new sediments that can be washed away on lithalsa slopes by surface runoff (Shilts, 1978; French, 2007) (Figure 3.9a). On the Sheldrake river catchment, where 1100 lithalsas covering an area of 3 km2 were inventoried in 2009

(Jolivel and Allard, 2013), this is shown by the milky colour of thermokarst ponds (Figure 3.9b). Rapid thaw at the base of the active layer also causes small active layer detachment failures (Harris and Lewkowicz, 1993) occurring on the side slopes of palsas and lithalsas and flowing into thermokarst ponds or streams (Jolivel and Allard, 2013). Larger slides occur along gullies and along river banks (Figure 3.9c); activity and contribution of gullies to fluvial load are confirmed by presence of small deltas at the confluence with streams or with the Sheldrake river (Figure 3.9d).

Rain Turbidity Thaw front DD ≥ 0 C

180 ⁰ 1600

160 1400

140 1200 120 1000 100 800 80 600 60 40 400 20 200

0 0 1-May 1-Jun 1-Jul 1-Aug 1-Sep 1-Oct

Figure 3.7 Turbidity (UTN), thaw front depth (cm), rain (mm/day) and cumulative degrees-day ≥ 0⁰C from 1 May to 1 October 2010.

Rain Turbidity DD ≥ 0 C

100 ⁰ 1200 90 1000 80

70 800

60 600 50 400 40 30 200 20 0 10 0 -200 1-May 1-Jun 1-Jul 1-Aug 1-Sep 1-Oct

Figure 3.8 Turbidity (UTN), rain (mm/day) and cumulative degrees -day ≥ 0⁰C from 1 May to 1 October 2013.

The high turbidity (80 NTU) recorded from 8 to 12 June 2013 occurred simultaneously to a significant increase in air temperature (+15 °C). It is also correlated with the onset of the period were air temperature maintains continuously above 0°C, as shown by the increase of the slope of the degree-days curve (Figure 3.8). No change in precipitation and thus, in discharge was recorded. It is likely that this increase in air temperature caused a rapid deepening of the active layer and triggered active layer landslides along the river bank, as observed in spring 2010 on the river bank (Jolivel and Allard, 2013). These landslides are widespread on the Sheldrake River catchment and can be an important source of sediment input (Jolivel and Allard 2013).

The maximum of the 2010 turbidity period has been reached on 10 July following several days of light rains (5/07 to 7/07: 16 mm; 9/07 to 13/07: 21 mm) (Figure 3.3). In 2013, a first peak of turbidity occurred on 17 July, after two days of light rain (16/07: 5 mm; 17/07: 4 mm). Then, the turbidity reached its maximum on 29 July, after 3 days of rain (27/07: 7 mm; 28/07: 13 mm; 29/07: 5 mm). These rain events probably increased the surface runoff and facilitated the washout of the thaw-released sediments. In both years, at the end of July/beginning of August, turbidity rapidly came back to lower values. A possible explanation is that surface, sediment-releasing and periglacial erosion processes slow down as the thaw front nearly stabilizes at the base of the active layer over permafrost areas, and when soils are finally thawed in the permafrost free areas of the catchment.

In the Arctic, hot summer temperature followed by moderate rains can trigger earth flows, active layer detachments slides and gully erosion, delivering considerable volume of sediment (Lamoureux, 2000; Lamoureux and Lafrenière, 2009, Lewis et al., 2012). However, a pulse of turbidity generally occurs within two or three days (Lewis et al., 2005; Lamoureux and Lafrenière, 2009). The 24 days period in 2010 and the 19 days period of high turbidity recorded in 2013 in the Sheldrake River catchment associated with soil thaw and permafrost processes such as thermokarst and frostboil activation are the first of the kind reported in the literature.

During these mid-summer periods of high turbidity, 53% in 2010 and 25% in 2013 of the total suspended sediment load was transported. More investigations reporting similar events are necessary to better understand the timing, the mechanisms and the impacts of such sediment releases.

3.7.3 Yearly measured rates of transport vs assessment of mass lost by thermokarst

On the Sheldrake river catchment, between 1957 and 2009, landslides associated with permafrost degradation have released roughly 48,505 t of sediments; that is a yearly average yield of 933 t.yr-1 of sediments (Jolivel and Allard, 2013). The yearly lost mass by thermokarst between 1957 and 2009 represents a potential sediment yield 6 and 9 times higher than the 2010 and 2013 suspended sediment released as estimated from river gauging (Table 3.4). Several explanations can be given for this discrepancy. First, a fraction

of the eroded sediments can be trapped in the catchment at low energy locations (river bed, streams, gullies, water tracks, ponds, lakes and topographic depressions). This is shown for instance by bottom sediments in thermokarst ponds (Bouchard et al., 2011; Bouchard et al., 2012). Second, the Sheldrake River catchment is now at a very advanced stage of permafrost degradation (Jolivel and Allard, 2013). Over the last 50 years permafrost decay and fragmentation have increased hydrological connectivity, enhancing sediment inputs in the drainage system by a factor of 1.6 (Jolivel and Allard, 2013). However, the part of sediment coming from this increased connectivity remains low in the overall budget, i.e. ~ 1% (Table 3.4).

River export CTP release Contribution Landslides Contribution (t) (t) (%) release (t) (%) TSS (2010) 160 1.01 / 1.46 0.7 / 1.1 932 583 TSS (2013) 105 1.01 / 1.47 0.7 / 1.2 932 888

Table 3.4 Sediment released by thermokarst in the total fluvial exports.

It is likely that thermokarst processes are now more active and so have more significant impacts on sediment release in more northern areas, such as the Nastapoka River basin, 50 km up north, where much more permafrost still remains in the catchment. Finally, evacuation of thermokart material mostly occurs during rain events that increase flow velocity (Toniolo et al., 2009). As a consequence, a large part of stored erosion products from the past years should have already been evacuated by the river flow. Therefore, the sediment loads measured in 2010 and 2013 are likely much less than during the five previous decades when thermokarst was active at a larger scale than nowadays.

3.8 Conclusion

The Sheldrake River is a small river representative of the eastern coast of Hudson Bay. Its hydrological regime is closely linked to annual air temperature variations and summer

81 precipitations. Its period of maximum discharge typically occurs during the breakup and snowmelt periods. The summer flow depends on rainfalls, which show good correlation with discharge and variations of sediment concentrations. However, the study reveals that warm mid-summer temperatures can lead to sustained high turbidity even when the river discharge remains low. Indeed, in 2010 and 2013, the period of high turbidity lasted for three weeks and occurred in July when warm summer air temperatures transfer heat into the active layer and provoke the rapid penetration of the thaw front to about 85-90% of the maximum seasonal depth. This is likely a time of the year when frostboils on permafrost and active layer detachment failures on the slopes of palsas and lithalsas are the most active, releasing fine-grained sediments that are rainwashed at the soil surface and transported to the stream system. The study shows that maximum sediment flow in a thermokarst region may be primarily thermally-driven rather than hydrologically-driven.

However, the advanced state of permafrost decay in the catchment suggests that fluvial sediment exports were much higher during the past decades. In a context of global warming and generalisation of permafrost degradation, the Sheldrake River catchment can reflect future processes that will occur at the scale of the circum-polar region. This kind of summer high turbidity event could have significant impacts on downstream aquatic and marine ecosystems.

3.9 Acknowledgements

This work received financial support from grants to M. Allard from ArcticNet and the Natural Science and Engineering Research Council of Canada. The Centre d’études nordiques of Université Laval provided important logistical support. We thank Denis Sarrazin for the installation and maintenance of the gauging station and Marc-André Ducharme for field assistance. The suggestions of Mickael Lemay greatly improve the manuscript. We are also grateful to the Inuit community of Umiujaq for its generous hospitality.

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CHAPITRE 4 - Morphostratigraphy and recent sedimentation in Nastapoka Sound, Eastern Coast of Hudson Bay

4.1 Résumé

Le Passage de Nastapoka, situé entre les côtes du Québec et l’archipel des îles Nastapoka, dans l’est de la baie d’Hudson, possède un relief subaquatique accidenté composé de crêtes assymétriques et de bassins profonds, favorables aux accumulations sédimentaires. La zone d’étude se situe au large d’une région où la dégradation du pergélisol et l’activité thermokarstique se sont accélérées au cours du 20ème siècle. Vingt-cinq carottes de sédiments marins ont été prélevées en avril 2009. Six d’entres elles ont fait l’objet d’analyses physiques et chronologiques; parmi elles, trois carottes furent finalement soumises à des analyses élémentaires (C, N, C/N) et isotopiques (δ13C, δ15N) afin de déterminer les sources de la matière organique sédimentaire. L’étude montre que les processus de sédimentation sont complexes et largement influencés par les courants de fonds. Ceci est confirmé par l’absence de sédiments colloïdaux dans plusieurs carottes et la présence de queues de comète et de chenaux d’érosion sur le fond marin. Les taux de sédimentation les plus élevés sont reportés près des côtes du Québec et de l’Île Gillies et les plus faibles dans les bassins profonds. Les analyses isotopiques et élémentaires du carbone et de l’azote montrent que la fraction de matière organique d’origine terrigène dans les sédiments de fond a augmenté de 30% depuis le Petit Âge Glaciaire et que ce phénomène s’est considérablement accéléré vers la fin du 20ème siècle. Bien que la dégradation du pergélisol soit probablement une source importante de cet apport accru en carbone terrigène, on ne peut l’isoler de d’autres sources potentielles liées aux changements environnementaux comme l’augmentation de la production primaire en milieu terrestre et marin.

4.2 Abstract

Nastapoka Sound, located between the coast of Québec and the Nastapoka Islands, in eastern Hudson Bay, has a complex subaqueous relief, formed of asymmetric ridges and

89 deep basins, favourable to sediment accumulations. The study area is situated offshore of a region where permafrost thaw and an increasing number of thermokarst features due to climate warming were reported over the late 20th Century. Twenty-five sediment cores were extracted in April 2009. Six of them were selected for physical and chronological analyses; among them, three were finally selected for elemental (C, N, OC/TN) and isotopic (δ13C, δ15N) analyses in order to identify sedimentary organic matter sources. The study shows that sedimentation processes are complex and primarily driven by bottom currents. This is confirmed by the absence of colloidal sediments in several cores and the presence of trailing-spits and erosion channels on the sea floor. The fastest sedimentation rates are found near the coast of Québec and of Gillies Island and the lowest ones occur in the deep depocenters. Carbon and nitrogen isotopic and elemental analyses reveal that the fraction of the sedimentary organic matter from terrestrial sources increased by 30% since the Little Ice Age. This trend has significantly accelerated toward the end of the 20th century. Permafrost decay is likely a major source for this terrestrial carbon. However, it cannot be distinguished from other potential sources that are also related to environmental changes such as increase in primary productivity both on land and at sea.

4.3 Introduction

The Hudson Bay bottom sediments are of a particular interest because they contain a record of the environmental changes that took place in the terrestrial environment of its surrounding catchment. Starting at deglaciation time, bottom morphology and sediments registered the catastrophic drainage of glacial Lake Agassiz-Ojibway into the North Atlantic (Lajeunesse and St-Onge, 2008). Several other studies have focused on the glacial/postglacial transition of the Hudson Bay through analyses of sediment cores and seismo-stratigraphy (Bilodeau et al., 1990; Gonthier et al., 1993; Hill et al., 1999; Lavoie et al., 2008).

Many studies using sediment cores, mostly in coastal sectors of the bay, successfully yielded paleo-environmental reconstructions (Jenner and Piper, 2002; Ladouceur, 2008; Haberzettl et al., 2010). More recently, studies have focused on atmospheric and fluvial processes affecting modern sedimentation in order to better understand the sediment

composition and sources (Hare et al., 2008; Kuzyk et al., 2008, 2009, 2010; Hülse et al., 2012). For instances, Kuzyk et al. (2008) showed evidences of temporal changes in the composition of terrigenous organic carbon in Hudson Bay sediments linked to the ice climate. Another study conducted by Kuzyk et al. (2010) focused on the relative importance of marine vs terrestrial organic matter (OM) input in Hudson Bay. Those studies, however, put the emphasis on the surface sediment layer and did not demonstrate any changes in the origin and accumulation rates of organic matter (OM).

During the late Holocene and particularly in recent decades, the coastal region on the east side of Hudson Bay has experienced significant variations in air temperatures, precipitations and sea-ice duration, which have affected environments by increasing their primary production and causing rapid permafrost thaw (Allard and Seguin, 1987; Payette et al., 2004; Bhiry et al., 2011; Chapter 2 and 3). However, sedimentary studies attempting to decipher a marine record of geomorphological and ecological changes in the watershed are almost absent, one exception being research reported near the mouth of Rivière de la Grande Baleine by Hülse et al. (2012).

This study was conducted in the Nastapoka Sound, offshore of a region where permafrost thaw and an increasing number of thermokarst features due to climate warming were reported over the late 20th Century (Marchildon, 2007; Larouche, 2010; Chapter 2). As the region was covered by vast expanses of frozen peatlands (palsa fields and plateaus) and by permafrost in fine-grained post-glacial marine sediments, it could be expected that the extensive decay of permafrost could have been a source of organic carbon and sediments that would have ended up being deposited in the marine environment, thus providing some record of climate-driven environmental changes. Therefore the objectives of this study are: (1) to describe and document the recent morpho-sedimentary dynamics through sediment accumulation rates and sedimentary features on the sea-floor of Nastapoka Sound, (2) to characterize surface sediments with an emphasis on identifying OM sources and (3) to identify possible changes in the sources of recent and modern sedimentary OM. The hypothesis tested is that the warming following the Little Ice Age (LIA) and its resulting impacts on permafrost decay have influenced transfers of OM and sediments from the terrestrial environment into the waters of Nastapoka Sound.

4.4 Study area

Hudson Bay is a fairly shallow southern continental extension of the Arctic Ocean which receives 30% of the total Canadian river flow over a catchment of 3.1 millions of km2 (Prinsenberg, 1986; Kuzyk et al., 2008). Marine water inflows coming from Arctic Ocean penetrate the bay via the northwest and leave it by the northeast through Hudson Strait, after experiencing a counter-clockwise water circulation (Prinsenberg, 1986).

The study area is located off the mouth of Sheldrake river, on the eastern coast of the bay, in Nastapoka Sound, 5 km north of the village of Umiujaq (Figure 4.1). This 4-6 km wide basin is located between the coast of Québec and the Nastapoka Islands; it sits between two sets of late Proterozoic volcano-sedimentary cuestas, with steep cliffs facing east and gentle slopes plunging west under the sea. It straggles parallel to the circular arch of the bay from Le Goulet of Lac Guillaume-Delisle to 170 km up North. This semi-closed sedimentary basin played an important role during the last deglaciation by trapping glacial, glaciofluvial and postglacial marine sediments (Lavoie et al., 2008).

Figure 4.1 Bathymetric map and location of the coring sites. Bathymetric data south of the unmapped areas come from Girard Thomas (2009).

Typically, three marine sedimentary units reflect the glacial and postglacial history of the region (Bilodeau et al., 1990; Gonthier et al., 1993; Hill et al., 1999; Lavoie et al., 2002, 2008): (1) a subaqueous ice-contact and glaciomarine unit deposited by the flowing ice sheet in contact with the Tyrrell Sea; (2) a paraglacial unit deposited during the transition

93 between glaciomarine to modern conditions by the reworking and resedimentation of emerged fluvioglacial sediment; (3) a Late Holocene to modern unit associated to fluvio- deltaic and marine processes. At the scale of Hudson Bay, the postglacial unit is composed of sediment derived from erosion of glacigenic sediments (120×106 t/yr), coastal erosion (18×106 t/yr), fluvial inputs (10.2×106 t/yr) and atmospheric deposition (0.74 ×106 t/yr) (Hare et al., 2008). Sea-ice rafting is another source of sediment but its input contribution is hard to quantify (Haberzettl et al., 2010). Acoustic profiles revealed a sediment thickness of 10 to 30 m and a maximum of 70 m in deep glacial troughs (Girard Thomas, 2008; Lavoie et al., 2008). The seafloor is very similar to the emerged topography with cuestas and glacial valleys and troughs oriented east-west (Lavoie, 2006, Girard Thomas, 2008).

In Nastapoka Sound, strong surface currents flow from South to North with a velocity reaching 15-20 cm.s-1 (Saucier et al., 2004). Bottom currents are supposed to be slower and reversed southward (Saucier et al., 2004), but their dynamics remain largely undocumented despite their influence on the postglacial sedimentation (Girard Thomas, 2008). The mean tidal range is more or less 1 m (Lavoie et al., 2002).

The area experiences a subarctic climate with cold winters (-24°C in January) and cool summers (10°C in July). As a result, Nastapoka Sound, similarly to the whole Hudson Bay, is totally frozen from mid-December to mid-June. However, recent global warming induces a later freeze-up and an earlier break-up (Gagnon and Gough, 2005).

In the region, the most important rivers flowing in the bay are the Nastapoka River (30 km north of the study area), the Little Whale River (70 km south) and the Great Whale River (170 km south) with an average yearly discharge of respectively 7.86, 3.74 and 19.77 km3.yr-1 (Déry et al., 2005). The 780 km2 estuary of Lac Guillaume-Delisle receives fresh water from four principal rivers and flows into Hudson Bay by a narrow and shallow channel named le Goulet, 50 km south of the study area. Annual discharge of this freshwater system is 4.49 km3.yr-1 (Déry et al., 2005). In Northern Canada, river discharge has increased since the early 1990’s (Déry et al., 2009). Large scale teleconnections, such as the Arctic Oscillations, are the main factors influencing the variability and trend of Canadian high latitude freshwater discharge (Déry and Wood, 2005).

The Sheldrake River drains a catchment covered by discontinuous permafrost that has been widely impacted by thermokarst over many decades (Chapter 2). In that river catchment, permafrost has decayed by 21% between 1957 and 2009, thus feeding a significant release of sediments and carbon out of the fluvial system (Chapter 2). This small river exported 160 t of sediments in 2010 and 105 t of sediments in 2013. 58% of this volume has been evacuated during a 24 days event of extreme turbidity (July) provoked by thermokarst and periglacial processes (Chapter 2).

4.5 Methods

4.5.1 Bathymetry and recovery of cores

Bathymetric surveys were performed in July 2008 from the MV Katherine-Anne with a Raymarine DSM300 digital sounder module (200 kHz) connected to a Raymarine C80 series display with NMEA output and coupled with an Edgetech 4100P sidescan sonar (100 kHz). Sixteen east/west or west/east (longitudinal) and 7 north/south or south/north (latitudinal) track lines were surveyed. Distance between two longitudinal lines was ̴ 250 m and distance between 2 latitudinal lines was ̴ 400 m. The goal was to map seabed morphology and to localise deeper areas which are likely to be efficient sediment traps for recent deposition from rivers of the eastern coast of Hudson Bay. The mapping starts at a distance ~1.5 km parallel to the coast of Québec until ~ 200 m from the shore of the Gillies Island. South of 56°38, depth data from Girard Thomas (2009) were used to complete the map. The map was realised with the ArcGis software by ordinary kriging, with 2028 selected depth points. Sidescan images were visualised thanks to the Discover 4100 software. Bathymetric data from Girard Thomas come from a PDF files. Girard Thomas used a Knudsen 320 B/P sonar (200 kHz) coupled with a GPS. Coupling was done by simple georeferenced superimposition with the ArcGis Software.

The coring sites were selected from the bathymetric map. Twenty-five short sediment cores were extracted from the ice cover in April 2009 with a K-B ® gravity corer (16 kg). The valve mechanism allowed minimal frontal wave effect and the water/bottom interface was well preserved. Cores are 20 to 40 cm long and have been extracted in depths between 52

95 and 82 m. Sediment cores were stored at 4°C until laboratory analyses. For this study, after a visual analysis of all the cores (Table 1), six of them were selected according to their location (03, 05, 10, 13, 23 and 24): they are located in the deeper areas and along a transect off the mouth of the Sheldrake River (Figure 4.1).

Core Longitude Latitude Depth Core Sediment texture 1 Sedimentary structures 1 IRD2 (m) lenght (cm) 1 76°33'5" W 56°38'13" N 52 7 coarse sand no yes 2 " " " 12 coarse sand no yes 3 76°34'35" W 56°37'49" N 70 39 very coarse silt to very fine sand no yes

4 76°36'0" W 56°39'25" N 80 34 massive silt (a) and sand (b) change in sedimentation from fine (a) to yes coarse (b) up-core (10 cm) 5 76°36'1" W 56°39'13" N 82 39 coarse silt to fine sand an erosion contact (10 cm) separates yes convolutes (downcore) and stratified sediment (up-core)

6 76°36'25" W 56°37'37" N 71 31 massive silt (a) and sand (b) sand bed (11 cm); convolutes (b) (19 cm yes downcore); change in sedimentation from fine (a) to coarse (b) (11cm) 7 76°36'2" W 56°37'59" N 62 36 massive silt (a) and sand (b) probable erosion contact (15 cm) associated yes with a change in sedimentation from fine (a) to coarse (b) up-core 8 76°34'59" W 56°38'52" N 65 27 silty sand 3 sand beds (10, 16, 22 cm) yes 10 76°34'0" W 56°40'29" N 73 33 coarse silt to fine sand 2 erosion contacts (20 and 28 cm) yes

11 76°35'0" W 56°40'24" N 74 31 silt no yes 12 76°35'47" W 56°40'13" N 77 30 silt no yes 13 76°35'1" W 56°40'0" N 72 38 very coarse silt no no 14 76°34'12" W 56°39'33" N 69 25 silt no yes 15 76°33'48" W 56°39'12" N 68 15 silt erosion contact (5 cm) with truncated yes convolutes 16 76°33'12" W 56°38'54" N 68 19 silt, sand and gravel no yes 17 76°33'6" W 56°38'31" N 64 15 silt, sand and gravel no yes 18 76°33'36" W 56°38'3" N 62 11 silt and sand no yes 19 76°34'45" W 56°38'13" N 65 34 massive silt (a) and sand (b) change in sedimentation from fine (a) to yes coarse (b) (13 cm) 20 76°36'9" W 56°38'47" N 82 38 massive silt (a) and sand (b) change in sedimentation from fine (a) to yes coarse (b) (23 cm) 21 76°35'26" W 56°37'46" N 70 31 silt sand bed (18 cm) yes 22 76°34'23" W 56°37'47" N 63 20 silt and sand no yes 23 76°33'33" W 56°37'48" N 61 27 coarse silt to fine sand no no 24 76°36'12" W 56°37'52" N 64 26 coarse silt to medium sand 3 sand beds (10, 14, 20 cm) yes 25 76°34'46" W 56°38'44" N 65 40 massive silt (a) and sand and gravel probable erosion contact (7 cm) with yes (b) change in sedimentation from fine (a) to very coarse (b) up-core 26 76°34'12" W 56°37'24" N 60 27 silt and sand contact erosion (7 cm); sand bed (22 cm) yes

4.5.2 Laboratory methods

All cores were first analyzed through CT-Scan (Institut National de la Recherche Scientifique, Quebec City) with a 1 mm resolution in order to visualize potential sedimentary structures, shells and ice-rafted debris (Orsi et al., 1994). CT-scanning also allows for imaging of density contrast showed by gray scale values expressed in CT- numbers (e.g. St-Onge and Long, 2009). CT-numbers can give indications about changes in mineralogy and organic matter content (e.g. Crémer et al., 2002; St-Onge et al., 2007; St- Onge and Long, 2009). We also used the CT-number as a proxy of density, since it has a higher resolution than by the method of gamma ray attenuation (Fortin et al., 2012). The

CT-number profiles represent a tomogram. The cores were then split and run for magnetic susceptibility (k) analyses using a GEOTEK Multi Sensor Core Logger (MSCL) at 0,5 cm intervals (Institut des Sciences de la Mer de Rimouski).

Grain size analyses were performed at 1 or 2 cm interval for each of the six cores. Carbonates and organic matter were removed using HCl and loss-on-ignition treatments. In order to exclude ice-rafted material, the fine fraction was separated from the coarse fraction with a 500 µm sieve. Only the finest fraction was kept and analyzed with a Horiba laser refractometer, after disaggregating in an ultrasonic bath. The results of at least three runs were averaged. Grain size measurements were calculated using Gradistat 8.0 (Blott and Pye, 2001). Mean grain sizes are used for interpretation and discussion.

4.5.3 Dating

210Pb and 137Cs analyses were performed at the radiochronology laboratory of Université Laval by gamma ray counting. Recent sedimentation rates were estimated from profiles of radioactive decay of 210Pb in association with the nuclear fallout 137Cs. Subsamples were extracted downcore at 1 cm interval. The “constant flux-constant sedimentation” model used here allowed estimating sedimentation rates in cm.yr-1 according to the slope of the ln- 210 linear regression of Pbin excess activity with depth.

The low level of 210Pb and 137Cs activity, close to detector background in some cores 210 (especially cores 05, 23 and 24), is a source of uncertainties. The activity of Pbin excess was measured graphically from the radioactive decay 210Pb profiles. The slope is calculated 210 from the ln Pbin excess versus depth graph.

Deeper samples in cores 13, 23 and 24 contained traces of 137Cs, preventing use of first entry of 137Cs in 1953 (date of first input in the environment). Thus, as usually done (Klaminder et al., 2012), sediment rates from 137Cs were calculated from the 1963 (year of fallout maximum) peak concentration as follows:

S= (Cmax-Lb)/(T-1963)

Where S is sediment accumulation rate (cm/yr); Cmax is the maximum concentration depth 137 (cm) of Cs; Lb is the mixing layer thickness and T the year of sample collection.

In addition, three shell fragments and two bulk clay samples were submitted to AMS 14C dating. 14C ages were calibrated using the CALIB 6.1 software (Stuiver and Reimer, 1993). A ΔR value of 180 (+/- 40) years was used for the reservoir effect correction (Barber et al., 1999; McNeely et al., 2006; Lavoie et al., 2012). Calibrated dates presented here are the median ages with ± 2 σ errors.

4.5.4 C, N, δ13C and δ15N analysis

C, N, δ13C and δ15N analysis were performed at G.G. Hatch Isotope Laboratories, University of Ottawa, on cores 03, 13 and 22. Each sample was divided into three subsamples. The first subsample was acidified twice with 10% HCl and then washed three times (dried and grounded) to remove carbonates. This subsample was used to determine total C (TC) and total N (TN) with a Carlo-Erba elemental analyser. The second subsample was used to determine residual N and C (which is considered to represent organic carbon (OC) and to analyse δ13C. The third one was used for δ15N. The second and the third subsamples were previously dried, grounded and sieved. OC/TN weight ratio is converted into molar ratio by multiplying by 14/12 (Lamb et al., 2006).

4.5.6 Historical changes in the fraction of terrestrial organic matter

To estimate the historical changes in the fraction of terrestrial organic matter (TOM) in marine sediment of the Nastapoka Sound, a two end member mixing model proposed by Calder and Parker (1968) was applied:

13 13 13 13 %TOM= (δ C marine - δ C Cobs) / (δ C marine – δ C terrigenous) × 100%

13 13 Where δ C marine and δ C terrigenous are respectively the marine and the terrestrial end 13 member values for Hudson Bay (Kuzyk et al., 2010), i.e. -20.8‰ and -28.2‰; and δ C obs is the measured δ13C of a given sample.

Sedimentation rates from 210Pb profiles were extrapolated until 1750, i.e. in the second half 210 of the LIA. In this study, the dating exceeds the maximum depth of Pbin excess, implying uncertainties. However, when sediments are homogenous in density and grain size, as in this study, it is common to date series up to 250 years (e.g. Liu et al., 2007).

4.6 Results

4.6.1 Sea floor morphology and distribution of surface sediments

Sidescan and cores revealed a complex surface sediment distribution. In general, fine sediments are located in the deepest areas or basins, while coarser material is concentrated near the coast, near Gillies Island and at the bottom of slopes. Ice rafted gravel and sand are dispersed over the sea floor. Bedrock outcrops appear as knobs. Bathymetric data and cores do not allow for description of the subaqueous relief and the sediment distribution in the shallow nearshore zone. However, at low tide, sedimentary deposition can be observed at the mouth of the Sheldrake River.

The two asymmetric east-west ridges (small cuestas ridges) in the center of the area are characteristic of the continental and marine regional landscape. Elevation between the top and the foot of the escarpment of the two ridges is ~ 40 m, while the dip slope extends over 1000 to 2500 m. Deep areas are situated along Gillies Island and in the northern part of the studied area. Depth reaches 85 m at the foot of the escarpment edge of the cuesta forming Gillies Island. Side scan imagery shows presence of trailing-spits and erosion channels at the base of ridge escarpments (Figure 4.2).

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Figure 4.2 Erosional landforms on the sea floor. (A) Trailing-spits (encircled); (B) erosion channel (arrow).

4.6.2 Chronology

210 137 210 Pb and Cs data and profiles are shown in Figure 4.3. According to profiles of Pbin excess, sedimentation rates vary between 0.02 and 0.13 cm/yr. The lowest accumulation rates occurred in the deepest central basin (0.02-0.06 cm/yr), while highest sedimentation rates are recorded near the coast (0.09 cm/yr) and near Gillies Island (0.13 cm/yr). A thin surface mixing layer is present in cores 23 and 24.

Radiocarbon dating on shell yielded ages between 1213 and 1757 years cal BP, while 14C dating on bulk sediment yielded ages of 2136 and 6394 years cal BP (Table 4.2).

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Core/Sample Sediment depth (cm) sample description 14C age (+/-) Age cal BP (2Ʊ range) Lab no. Core 23a 14 Unidentified shell fragments 1845 (15) 1213 (1123-1304) ULA-1593 Core 24 21.5 Unidentified shell fragments 2320 (15) 1726 (1605-1847) ULA-1589 Core 23b 16 Unidentified shell fragments 2350 (15) 1757 (1639-1875) ULA-1591 Core 03a 20.5 bulk sediment 2650 (20 2136 (1995-2278) ULA-1764 Core 03b 30 bulk sediment 6165 (20) 6394 (6287-6502) ULA-1765

Table 4.2 14C dates

4.6.3 Lithology and grain size

In the six cores, sediments are visually similar (Figures 4.4 and 4.5). They consist of compact medium silt to medium sand with rare shell fragments, traces of bioturbation, some black reduction spots, and scattered drop sand and gravel. Mean grain size ranges from 27 to 64 μm. The colloidal portion is null or negligible. The grain-size distributions are generally bi-modal and very-poorly sorted (Figure 4.6). Three erosion contacts are recorded in cores 05 and 10. They are shown by truncated beds on CT-scan images and by changes in the CT-number, mean grain size and k.

Core 05 registered soft sediment deformations from ~12 cm downward (Figure 4.4). Convolutes, in the form of sand waves and wavy and non parallel stratifications, are observed on the CT-scan image and by variations in k and CT-number. In cores 13, 24 and 05, mean grain size tends to increase toward the sediment surface, while cores 03, 10 and 23 show no significant trends.

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CORE 05

CORE 10

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CORE 24

Figure 4.4 CT-scan image, mean grain size (µm), magnetic susceptibility (k) and CT- number of cores 05, 10 and 24. A zoom of the CT-scan image of core 05 highlights the convolutes and the erosion contact. Light gray bands show erosion contacts in cores 05 and 10; darker gray bands highlight sand beds in core 24. The black star indicates 1900 AD inferred from 210Pb profiles.

CORE 03

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CORE 13

CORE 23

Figure 4.5 Physical, elemental and isotopic profiles of cores 03, 13 and 23. The black star indicates 1900 AD inferred from 210Pb profiles.

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5,0

4,0

3,0 Class (%) 2,0

1,0

0,0 1 10 100 1000 Particle Diameter (mm)

Figure 4.6 Average grain size frequency for core 13.

4.6.4 Elemental and isotopic composition of sedimentary OM

Sediments have very low TN and OC contents of ≤ 0.15% and ≤ 0.8% respectively. In general, values are constant or increase slightly toward the surface, except TN in core 23 which shows an opposite pattern. C/N molar ratios vary between 1 and 11. δ13C values range from -23.3 to -25.1 ‰, while δ15N values range from 6.5 to 9.2 ‰. The three cores registered a general decrease in δ13C and in δ15N toward the surface, except core 03 that does not show a significant vertical trend in δ15N. Core 23 registered a clear steady increase in the C/N molar ratio toward the surface, whereas core 03 and 13 profiles are rather chaotic. On the other hand, in all cores, the surface layer (2-3 cm) is marked by a more pronounced increase in C/N while δ13C and δ15N decrease significantly (Figure 4.5).

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4.7 Discussion

4.7.1 Sedimentation rates and chronology

Radiocarbon dates of three shell fragments on cores 23 (2 samples) and 24 yielded ages of 1213, 1657 and 1726 years cal BP, inferring sedimentation rates of ~0.01cm/yr (Table 4.2). No erosion surface is visible in these cores. 210Pb chronology shows sedimentation rates of 0.09 (core 23) and 0.13 cm/yr (core 24). Sedimentation rates off the Goulet at the outlet of the Lac Guillaume Delisle (Haberzettl et al., 2010) as well as those off the mouth of the Great Whale River are of the same order of magnitude (Jenner and Piper, 2002; Kuzyk et al., 2008; Hülse and Bentley, 2012). Therefore, it is likely that the three shell fragments provide inexact ages either because they have been redeposited or because the ΔR was underestimated.

Two dates on carbon from bulk sediments were also obtained on core 3. Results show ages of 2136 years cal BP and 6394 years cal BP (Table 4.2). As for shell dating, those ages appear considerably older than expected with the 210Pb chronology and other studies in the region. Numerous peatlands affected by riverbank and coastal erosion are a potential source of old carbon in the coastal marine environment (Guo et al., 2004). It is thus likely that recycled older carbon was introduced in the sedimentary system, as was also observed in bottom sediments of thermokarst ponds near Kuujjuarrapik (Bouchard et al., 2012).

210 For the six cores, ln Pbin excess is well correlated with depth (Figure 4.3), despite the presence of a thin surface mixing layer in cores 23 and 24. This linear pattern reflects constant sediment rates, as observed off the Great Whale River mouth (Hülse and Bentley, 2012). Conservation of bioturbation overprint confirms slow sediment accretion rates that allow for biological activity (Bentley et al., 2006). Low sedimentation rates also imply large time intervals per centimeter (1 cm = 8 to 50 year of sedimentation). This may explain the bimodal and the very-poorly sorted grain size.

The 210Pb profiles are validated by 137Cs (Smith, 2001). For cores 03, 13 and 24, the depth of maximum concentration of 137Cs confirmed the 210Pb chronology. On the other hand, the 137Cs maximum concentration is deeper than expected with profiles of 210Pb in cores 10 and

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23 (and to a lesser extent in core 24). This can be linked to the downward diffusion of 137Cs as a result of surface sediment mixing (Robbins et al., 1978) and of diffusion in pore water (Klaminder et al., 2012). On the opposite, in core 05, the 137Cs maximum concentration is shallower than expected. The presence of 137Cs near the surface sediment and the relative low amplitude of the peak of 137Cs can be associated with a deferred supply of 137Cs occurring when freshwater sediments are transferred to the marine environment (Oughton et al., 1997; Klaminder et al., 2012). This process may be enhanced by remobilization of emerged sediments in a context of isostatic rebound. Finally, the 1 cm resolution subsampling and the low sedimentation rate imply a non-negligible margin of error about the maximum concentration depth of 137Cs.

However, if the depth of 137Cs maximum concentration truly corresponds to 1963, the difference of sediment rates registered by the 210Pb and the 137Cs methods would reflect a change in sedimentation rate as the two methods have a different time interval, i.e. ~150 years for the 210Pb and ~60 years for the 137Cs. Cores 10 and 23 are the closest to the coast and directly off the mouths of two small rivers, while core 05 was extracted in the deeper offshore area. This interval between the two methods of dating was also observed off the Great Whale River mouth by Hülse and Bentley (2012), where it is interpreted as a consequence of a more energetic environment which transported fluvial fine sediments further offshore. However, no grain size proxy confirms this hypothesis.

The mean local sediment accumulation rate (0.07 cm/yr) is in accordance with the only available 14C chronology in the area (outlet of Lac Guillaume Delisle, 0.09 cm/yr in Haberzettl et al., 2010). However, it is slightly lower than in central and western Hudson Bay (Kuzyk et al., 2008, 2009) and eastern Hudson Bay influenced by large rivers inputs, such as the Great Whale River and the Innuksuac River (Jenner and Piper, 2002; Kuzyk et al., 2008, 2009 ; Hülse, 2012). This supports the interpretation that the Nastapoka Sound is an energetic environment with strong currents (Saucier et al., 2004) and that sediment inputs from rivers of the south-eastern Hudson bay are either deposited in the proximal river mouth area or dispersed in northward direction in the counterclockwise current circulation and in the Hudson Bay system in general (Hülse and Bentley, 2012).

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In the studied area, sedimentation rates range from 0.02 to 0.13 cm/yr depending on location. Deep basins generally registered lower sedimentation rates than shallow coastal sites. However, there is no relationship between sediment rates and water depth, distance from coast or distance from the Sheldrake River mouth, as observed off the Great Whale River (Hülse and Bentley, 2012). This confirms that deposition processes are primary driven by local marine hydrodynamics and are not significantly affected by rivers located 70 km (the Little Whale River) and 170 km (the Great Whale River) to the south. In the proximal area, rivers have low discharge (Sheldrake River mean annual discharge: 10 m3/s, Chapter 3). Consequently, fluvial sediment inputs have no sufficient influence on local sediment accumulation rates to offset marine processes.

4.7.2 A complex sedimentary regime

4.7.2.1 A dynamic area

Previous seismo-stratigraphic and coring studies in the area revealed that the sediments below the surface are postglacial hemipelagic sediment from suspension under a low- energy sedimentary regime (Lavoie, 2002; Lavoie et al., 2008; Girard-Thomas, 2009). Subaqueous topographic depressions, such as at the foot of the Gillis Island, acted as traps for glacigenic, glaciomarine and postglacial sediments, as shown by the great thickness of the sediments deposits in these depocenters (Girard Thomas, 2009, Lavoie, 2006; Josenhans et al., 1988). However, lack of seismo-stratigraphic data does not allow for generalizing to the whole studied area. Indeed, strong bottom and wave erosive currents can lead to the exposition of glaciomarine sediment on the sea floor. This was observed in the Nastapoka Sound (Lavoie et al., 2008; Girard Thomas, 2009), in the Manitounouk Sound (Hill et al., 1999) and on most of the central Hudson Bay seafloor (Josenhans et al., 1988).

In the studied area, several lithological features and bottom landforms reflect the strong influence of bottom currents. This influence is shown by the quasi-absence of colloidal sediments in every core, erosion contacts in cores 05 and 10, presence of trailing-spits, erosion channels (Figure 4.2) and possible current megaripples on the seafloor (Girard Thomas, 2009). The uneven subaqueous relief seems to influence bottom currents, which

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increase in velocity along topographic sills and are channelled into erosion scours (Girard Thomas, 2009; Figure 4.2).

4.7.2.2 Disturbed sediment deposits: core 05

In core 05, reworked sediments, in the form of convolutes are evidenced in stratified sediments, with black OM rich beds (Figure 4.5). This suggests a recent and continuous sedimentation. However, truncated stratification reveals an erosion contact between the two facies. The very low sedimentation rates and the thin recent sediment layer exclude the hypothesis that the convolutes are recent load structures (<600 yrs).

The reworked unit has not been dated because it is too deep for 210Pb and 137Cs dating, no 14C datable material was found and no deeper seismic data were available. The extrapolation of sedimentation rates calculated from 210Pb profile (0.02 cm/yr) suggests that the contact between the two facies occurred by ~1400 AD, but the time interval lost during the erosion phase remains unknown.

This kind of convolute is likely made by gravity flows and liquefaction of soft sediments (Lajeunesse and Allard, 2002). Mass movements and landslide deposits have been identified in the Manitounouk and Nastapoka sounds in two types of sediments: recent marine (or deltaic) deposits and glaciomarine (or paraglacial) basin-fill deposits (Hill et al., 1999; Lavoie et al., 2002; Lavoie et al., 2008; Girard Thomas, 2009).

Mass movement deposits and landslide scars have been observed in recent marine sediments of the Nastapoka Sound and Lac Guillaume Delisle (Lavoie et al., 2002, 2008; Girard Thomas, 2009). Hudson Bay is not a seismic area (Adams and Halchuk, 2003) and the low sediment rates in the Nastapoka Sound limit these mass movements to coastal and shallow areas (Lavoie et al., 2008), on steep slopes of subaqueous cuestas (Girard Thomas, 2009), as well as on the front slope of deltas (Lavoie et al., 2002). Moreover, scars of translation slides observed by Girard Thomas (2009) do not bear internal disturbance as the slide occurred on a shear surface (i.e. here glacial polish on rock outcrops).

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The only seismic unit showing evidence of gravity flow, hummocky deposits and disturbed sediment deposits in the Nastapoka and Manitounouk sounds is a rapid basin infilling unit deposited during the transition between glaciomarine and marine conditions (Hill et al., 1999; Lavoie et al., 2008; Girard Thomas, 2009). The location of the coring site in a deep depot-center and the fact that this unit can be affected by erosion (Lavoie et al., 2002) would support the hypothesis that the disturbed sediment observed in core 05 belongs to this seismic unit.

Assuming this origin, a possible cause of slide triggering is the seismic activity associated with rapid isostatic uplift during or shortly after the disappearance of the Laurentide ice sheet from eastern Hudson Bay (Hill et al., 1999). During this period, gravity flows evolving into turbidity currents would have been common in the area (Hill et al., 1999, Lajeunesse and Allard, 2002, 2003).

The significance of the thin overlying stratified sediment is ambiguous. It implies a change in the bottom current and sedimentary regime, from erosion to sedimentation conditions. This transition would have occurred by ~1400 AD. Contemporary marine conditions have existed since 6500 BP, i.e. since the end of the glaciomarine phase (Bilodeau et al., 1990). Continuous fall of the relative sea level has likely influenced the water circulation and sediment transport in the sound, as a result of the progressive closure of the Manitounouk Sound further south (Allard and Tremblay, 1983). In the absence of seismic data and as the side scan imagery does not reveal any specific landforms, the only remaining explanation is a change of sedimentary environment. It is hypothesized that an erosion channel flowed in the center of this deep basin and was then filled. Such a filled channel has been observed by Girard Thomas (2009) on side scan imagery at the foot of a cuesta escarpment and at a similar depth. A change in bottom currents or an increase of sedimentary inputs would have stopped the activity of channel erosion and allow for sedimentation. This hypothesis is also supported by changes of sedimentation from fine to coarse material, separated by potential erosion surfaces, recorded in 6 other cores, located at the bottom of slopes, and particularly in cores 4 and 20 located in the same depocenter of core 5 (Figure 4.1).

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4.7.2.3 Bottom current deposits

Cat-scan image of core 24 shows sand beds characterised by an increase in mean grain size and CT-number and a decrease in k (Figure 4.4). This sequence of sand beds does not seem to be linked to periods of high discharge leading to hyperpycnal flow (Mulder et al., 2003) as it was not recorded in any other cores located near or in front of the Sheldrake river mouth (Table 4.1). First entry of 210Pb is deeper than the upper-most sand bed, and no erosion contact is visible, suggesting that the upper sand beds are recent (~1940 AD for the upper one). The relatively low resolution of our proxies compared to the thinness of the sand layers does not allow for identification of sediment grading associated with distal deposits of turbidity currents.

The bathymetric map shows that core 24 was extracted at the foot of a steep slope, on an elevation of the sea floor between two deeper basins. These narrow corridors between two topographic obstacles are favourable to increased bottom current velocity (Hill et al., 1999; Girard Thomas, 2009). However, this core registered the fastest sedimentation rate of the area, suggesting that bottom current deposits are a significant source of sediment in this particular area.

Four other cores, which were not physically analysed, have sand beds (Table 4.1). Three of them (cores 6, 8 and 21) are located in the same topographic environment, i.e. in submarine valleys or corridors (Figure 1). Thus, sand beds may be lags created by recurrent periods of highest energy. According to that interpretation, the gradually upward increasing mean grain size, also observed in cores 05 and 13 (Figures 4.4 and 4.5), could be related to a recent increase in energetic marine conditions in Nastapoka Sound, as proposed by Hülse and Bentley (2012) for the Manitounouk Sound.

4.7.3 Recent sedimentation: increase of terrigenous influence

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4.7.3.1 δ13C and C/N molar ratio

δ13C and C/N are efficient tracers for terrigeneous organic matter in Hudson Bay (Kuzyk et al., 2010), because they provide an estimate of the relative proportions of marine and terrestrial sources of OM in marine sediments (e.g. Naidu et al., 1993; Meyers, 1994; Thornton and McManus, 1994; Lamb et al., 2006). δ13C and C/N in surface sediment both decrease with increasing water depth, latitude and distance from the coast in Hudson Bay (Kuzyk et al., 2010) and in the Arctic Ocean (e.g. Naidu et al., 2000; Nagel et al., 2009). This is due to the progressive mixture of terrestrial and marine OM along a distance to shore gradient.

The isotopic signature of OM is preserved in sediments through long periods of time, even if organic matter may continue to be degraded (Meyers et al., 1994). However, decomposition of OM during transport, deposition and diagenesis can be a source of biochemical alteration of isotopic and elemental carbon properties (e.g. Thornton and McManus, 1994; Meyers, 1997; Lamb et al., 2006). In this study, no evidence of diagenesis, i.e. C/N and δ13C are not correlated with sedimentation rates (Kuzyk et al., 2010); no significant change of mean grain size that could biases δ13C and C/N molar ratios (Gearing et al., 1977; Meyers, 1994; Thornton and McManus, 1994; Meyers, 1997) was found in the vertical sequences. However, the low contents of N (≤ 0.15%) and OC (≤ 0.8%) can be a source of large uncertainties (Sampei and Matsumoto, 2001; Lamb et al., 2006; Lavoie et al., 2008) and do not allow for interpretations of downcore changes.

δ13C and C/N molar ratio in the surface sediments of the Nastapoka Sound show an influence of both marine productivity and terrestrial C3 plants (e.g. Lamb et al., 2006), as shown by Kuzyk et al. (2010) for the whole Hudson Bay. Moreover, δ13C versus C/N molar ratio in Nastapoka Sound are distinctive from the rest of Hudson Bay (> 25 km offshore) but similar to values found in Lac Guillaume Delisle (Figure 4.7). This indicates that Nastapoka Sound is an area that undergoes a significant terrestrial influence, i.e. similar to the estuarine conditions of Lac Guillaume Delisle. Nastapoka Islands likely act as a barrier by holding freshwater and nutrient input released by numerous rivers flowing in James Bay and eastern Hudson Bay (Hudon et al., 1996; Déry et al., 2005). Moreover, the counter

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clockwise water circulation and strong coastal currents may limit offshore dispersion of river input. Finally, although it is situated in the deepest area, core 03 recorded the absolute highest terrestrial signal, indicating that deep depocenters are efficient traps for terrigenous sediment.

Figure 4.7 δ13C vs C/N in the Nastapoka Sound (white squares) compared with values in Lac Guillaume Delisle (black triangles) and in the rest of Hudson Bay (black circle).

4.7.3.2 δ15N

In correlation with C/N molar ratio and δ13C, δ15N of sedimentary OM also provides an estimate of contributions of marine and terrestrial sources (e.g. Naidu et al., 2000; St-Onge and Hillaire-Marcel, 2001; Nagel et al., 2009). δ15N tends to decrease while moving away

115 from the coastline (e.g. Guo et al., 2004; Gaye et al., 2007; Nagel et al., 2009). In Hudson Bay surface sediments, there is a general spatial correspondence of δ15N with C/N and δ13C: highest δ15N values are found in the central and western Hudson Bay and lower values are recorded in the south and eastern Hudson Bay, showing a more relative terrestrial influence in the south of the Bay (Kuzyk et al., 2010). Our study confirms this trend: δ15N of surface sediments in the Nastapoka Sound is similar to values found in southern and eastern coastal Hudson Bay and lower than in sediments of the central Hudson Bay.

δ15N is sometimes used as an isotope signal from permafrost soils (Nagel et al., 2009). Indeed, low values of δ15N may indicate colder climate with tundra vegetation or a catchment where permafrost was more abundant than today (Voss et al., 2006; Nagel et al., 2009). Additionally, OM released from permafrost catchments is less degraded than in permafrost free catchments and δ15N stays low (Guo et al., 2004; Nagel et al., 2009). A general decrease of δ15N, as observed in cores 13 and 23, suggests an increase of permafrost thickness and spatial distribution and a rarefaction of vegetation cover, associate with colder conditions. However, recent changes in landscapes around south-eastern Hudson Bay registered the opposite trend (e.g. Payette et al., 2004; Bhiry et al., 2011).

This study indicates that the sedimentary OM of the Nastapoka Sound comes from a mix of terrestrial and marine sources (Kelly et al., 2005). However, the absence of correlation between δ15N and δ13C and between δ15N and C/N reveals that δ15N is not only influenced by the relative mixing of the marine/terrigenous sources. The increase of δ15N toward the surface in cores 13 and 23 also indicates that permafrost decay since the LIA had no significant impact on N from OM released in the Nastapoka Sound. These findings support the idea that river inputs are a relatively low source of N compared to upwelling of deep water in coastal regions of Hudson Bay (Kuzyk et al., 2010). δ15N in surface sediment is mostly influenced by the variability of the δ15N of the phytoplankton in surface water and post-production processes (Kuzyk et al., 2010). Consequently, δ15N appears as an inefficient proxy for studying change in the OM sources in Nastapoka Sound and as an indicator of change in permafrost extent over land.

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4.7.3.3 Historical changes in terrigenous OM contribution since the LIA

Using Equation 1 and extrapolated sediment rates, the historical proportion of TOM since the second part of the LIA has been represented for cores 03, 13 and 23 (Figure 4.8).

Core 03 Core 13 Core 23

% of terrestrial organic matter 30 40 50 60 30 40 50 60 30 40 50 60

2000

1950

1900

1850

1800

1750

Figure 4.8 Variation in the proportion of terrestrial organic matter in cores 03, 13 and 23 since 1750 AD. Arrows highlight common trends.

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In the three cores, minimum of TOM are recorded at the end of the LIA, in the early 19th century. During the LIA, terrain conditions were drier than today, which supposes lower river discharge, and so less TOM input in Hudson Bay. These drier conditions are shown by the low water level of lakes as a result of snow deficit until ~ 1750 AD and a higher frequency of forest fire in Northern Québec (Payette and Fillion, 1993; Payette et al., 2008). Moreover, an extended and longer ice cover and lower surface water temperatures of Hudson Bay during the LIA may have led to a lower marine primary production, inducing an increased influence from terrigenous sources. Also during LIA, the permafrost was aggrading in the region (Allard and Séguin, 1987) limiting thermokarst which can be a significant source of terrigenous inputs in fluvial systems (Chapter 2).

The terrigenous fraction increased significantly from the first part of the 19th century (cores 03 and 13) and from 1850 AD (core 23). In core 13, this pattern coincides with an increase in mean grain size and k and a decrease in the CT-number. Since ~1850 AD, mean annual air temperatures, precipitations and river water level gradually increased in Northern Québec (Payette and Delwaide, 1991, 2000), which likely resulted in a rise of fluvial inputs into Hudson Bay. This period of higher discharge corresponds with low δ13C in coastal surface sediment (Gaye et al., 2007). The permafrost also started to decay at the end of the 19th century (Payette and Delwaide, 2000). As a result, thermokarst showed a growing activity. Permafrost thawing was then continuous during the 20th century (Payette et al., 2004). A significant rise in summer surface water temperature and salinity and a decrease in the duration of the ice cover, leading to an increase of marine productivity were also observed at the end of the 19th century (Ladouceur, 2008). This trend does not clearly appear in the TOM proportion neither in the indicators of production (OC, TN), but may be reflected by an increase in δ15N in the three cores around 1900.

The three cores revealed a decrease of TOM proportion in 1940-1950 AD with a minimum reached around 1960 AD. The time interval associated with each 1 cm sample corresponds to 11 to 17 years. Because of this poor time resolution, these minimums likely reflect a multi-year trend. Unfortunately, no discharge data are available in the region for this period. However, the end of the 1950’s and the 1960’s correspond to a worldwide decrease in air temperature (Jones et al., 1999) and more specifically to a decrease in snow

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precipitations in the region (Payette and Delwaide, 1991; Payette et al., 2004). A thinner snow pack would have resulted in lower melt water during snowmelt, reducing the intensity of the spring freshet during which much of sediment and nutrient are exported into Hudson Bay (Hudon et al., 1996).

The three cores registered an important increase of TOM influence in the recent decades. This increasing trend is likely a response of the recent increase in air temperature recorded in North-Eastern Canada since the early 1990’s (Chouinard et al., 2007), correlated with a general increase in river discharge (Déry et al., 2009). This recent trend is also confirmed by the decrease in δ15N and the rise in C/N in the first centimetres toward the surface of each core. The recent climate change may affect the proportion of TOM in the surface sediment of the Nastapoka Sound in different ways. Permafrost has degraded since the end of the LIA; however, its rate of decay has considerably increased since the beginning of the 1990’s (Payette et al., 2004). For example, the hydrological connectivity within the Sheldrake river catchment and permafrost degradation processes such as thermokarst and landslides, have significantly increased in number and extent, enhancing fluvial load released in Hudson Bay (Jolivel and Allard, 2013, Chapter 3 of this thesis). In fact, the closest core from the Sheldrake River mouth (core 23), shows the most obvious trend of an increasing TOM influence, i.e. the decreased of δ13C is supported by a continuous increase in C/N and decrease in δ15N (Figure 4.5). The constant mean grain size of core 23 suggests no change in Sheldrake River discharge and the increase of TOM proportion could be interpreted as a local signal reflecting permafrost decay in the Sheldrake River catchment. In fact, a previous study revealed that the turbidity peak of the Sheldrake river is more linked to activation of permafrost degradation processes than to peak of the river discharge (Chapter 3). Moreover, the accelerated rate of erosion of widespread peat deposits and palsas since the early 1990’s (Payette et al., 2004) can also be an important source of terrigenous organic matter in that coastal area (Brown et al., 2003).

Over the last four decades, Hudson Bay sea ice began to form later and break up earlier, particularly along the southern shore (Gagnon and Gough, 2005), losing an average 11.3% of its summer sea ice cover per decade (Tivy et al., 2011). This lengthening of the ice free period favours more energetic marine conditions and offshore transport of terrestrial matter

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(Hülse and Bentley, 2012). Due to the Hudson Bay’s counterclockwise circulation, it is likely that a part of the terrestrial organic mater released by large rivers south of the area (Great Whale River, Little Whale River) disperses along the eastern coast of Hudson Bay (Hülse and Bentley, 2012) and reaches our study area. In this context, it is possible that an increasing volume of this material settles in the Nastapoka Sound, even if no evidence is observed in mean grain size and in sedimentation rate. Later freeze-up and earlier breakup also imply that coastal areas are less protected from erosion by waves, especially during fall storms. Thus, coastal erosion could also potentially an important source of terrestrial organic carbon which can contribute to lower δ13C in the marine environment (Rachold et al, 2000). However, the coastline of the bay and the islands in the region is dominated by rocky ledges and gravel beaches. Fluvial inputs from river draining catchments affected by decaying peat rich permafrost, by other erosion processes and partially covered by wetlands are therefore the most probable source for the recent increase in TOM. For instance, a lignin study revealed that sedimentary terrigenous organic matter in southern Hudson Bay is composed from a mixture of angiosperm and gymnosperm and woody and non-woody material sources (Kuzyk et al., 2008). This is supposed to come from several sources such as erosion of peat deposits, river banks, organic surface soil and old marine and fluvial sediments.

Finally, development of soils and vegetation in this zone of transition can be another factor leading to an increase of terrestrial input, as shown by Miltner et al. (2005) and Lavoie et al. (2008) following the last deglaciation. Recent climate change has contributed to an increase of the shrub cover (Myers-Smith et al., 2011), particularly in coastal areas of Hudson Bay since the early 20th century (Laliberté and Payette, 2008). This is likely to increase the influence of C3 plants on organic matter composition released by rivers, lowering its δ13C. However, these changes seem too slow and too recent to be recorded in marine sediment. Moreover, modern plants debris is supposed to be retained near river mouth as a result of hydrodynamic sorting (Kuzyk et al., 2008).

The fraction of TOM in the surface sediment of the Nastapoka Sound varied trough time since the LIA. Proportion of terrigenous vs marine organic matter has increased by 30% (Shel03: +41%; Shel13: +19%; Shel23: +28%) since the end of the LIA. It is now ≥50% in

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the surface sediments of the Nastapoka Sound, which is higher than in offshore areas (≥25 km from the coastline) of southeastern Hudson Bay, where it varies between 15 and 35% (Kuzyk et al., 2010). The same general trend appears in the three cores with more or less amplitude and precision. Uncertainties about sedimentation rates deeper than allowed by 210Pb and the large time interval represented in each one centimetre subsample can explain this difference. The present study shows that variations in the proportion of TOM in recent marine sediments are linked to environmental changes resulting from climate warming.

4.8 Conclusion

Nastapoka Sound is an energetic marine environment. Subaqueous topography is complex and favours strong bottom currents, which control processes of sedimentation and erosion, as shown by trailing splits and erosion channels on the sea floor. Low sedimentation rates, erosion surfaces, lags, a thin postglacial sediment layer, changes in sediment texture and the absence of colloidal sediment in cores also reflect this influence. The velocity and the direction of bottom currents vary through time, likely in response to change in hydrodynamics due to the postglacial isostatic rebound.

Despite strong marine hydrodynamics, the surface sediments of Nastapoka Sound recorded changes in the source of organic sediment. Nastapoka Sound is an area of mixture between marine and terrestrial inputs and can be compared to an estuarine system. The downcore study of OM quality shows that TOM input is now superior to marine OM. Since the end of the LIA, contribution of TOM to marine sedimentation has increased by 30%. Despite some chronological and analytical uncertainties, it is clear that the coastal eastern Hudson Bay has experienced considerable changes in sediment input and quality since the LIA. The evolution of the terrestrial ecosystems affected by permafrost decay very likely took an important part in the observed marine changes.

4.9 Acknowledgments

This project was supported by grants from Arcticnet and the Natural Science and Engineering Research Council of Canada (discovery grant to M. Allard). The Centre

121 d’études nordiques of Université Laval provided important logistical support. We sincerely thank Bill Doidge for its availability and the sailing equipment. Thanks are due to Emmanuel L’Hérault, Maud Audet Morin and Carl Barrette for help in mapping, and Donald Cayer, Valérie Mathon Dufour and Catherine Falardeau-Marcoux for assistance in grain size studies. The suggestions of Mickael Lemay are appreciatively acknowledged. We are also grateful to the Inuit community of Umiujaq for its generous hospitality and support on the ice, particularly Joshua Sala and Peter Novalinga.

4.10 References

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Chapitre 5 - Conclusion générale

Cette étude a permis d’améliorer les connaissances sur les processus thermokarstiques et les taux de dégradation du pergélisol dans la zone discontinue du Nunavik, dans une région où le pergélisol est particulièrement riche en glace et où le réchauffement climatique est rapide. Les quantités de sédiments organiques et minéraux érodés et, donc, remis en circulation dans le géosystème ont été évaluées (objectif principal). Les résultats mettent en évidence le rythme soutenu de dégradation du pergélisol ainsi que les impacts environnementaux qui en découlent. L’approche systémique utilisée dans cette thèse a permis d’établir un lien entre les processus de dégradation du pergélisol et le transport en suspension du produit du thermokarst à l’échelle d’un bassin versant, sous forme d’entrées et de sorties.

La première partie de cette thèse (chapitre 2) montre que le paysage de la région d’étude s’est grandement modifié entre 1957 et 2009. L’étude confirme la dégradation très rapide du pergélisol sur la côte est de la baie d’Hudson déjà notée par d’autres chercheurs. Elle en précise les modalités et les taux, notamment en prenant en compte les mouvements de masse et l’érosion des berges, ce que n’avaient pas fait les recherches précédentes. Cette dégradation se caractérise par une diminution de la surface couverte par les lithalses, les palses, les plateaux palsiques et les plateaux de pergélisol, un doublement de la surface couverte par les mares de thermokarst et une augmentation du nombre de formes d’érosion actives comme les glissements de terrain (objectif 1). Ce chapitre décrit, pour la première fois, des formes d’érosion thermokarstique répandues en milieu de palses et lithalses : glissements de la couche active favorisés par l’érosion fluviale, ravins d’érosion entre deux plateaux de pergélisol (objectif 2). Il est montré que ces formes d'érosion thermokarstique sont une source potentielle de sédiment et de carbone organique pour le système fluvial (objectif 3). L’influence climatique de la baie d’Hudson est considérable et l’étude montre que la limite des arbres est une frontière prépondérante entre deux styles et deux vitesses de dégradation du pergélisol: le thermokarst est significativement plus intense dans la toundra forestière que dans la toundra arbustive. Ce changement de paysage et l’importance de cette limite biogéographique se reflète également dans l’augmentation de la connectivité hydrologique établie par la formation de ravins et de corridors entre les mares de

129 thermokarst, laquelle engendre une capacité accrue de transport des volumes de sédiment et de carbone organique rejetés dans le réseau de drainage (objectif 4). Toutefois, il est montré que le bassin versant de la rivière Sheldrake, malgré son activité thermokarstique intense, ne rejette pas davantage de sédiment ni de carbone organique dans le réseau fluvial qu’un bassin versant sans pergélisol.

Le troisième chapitre prend la suite du deuxième et se concentre sur les processus de remobilisation et de transport du matériel produit par l’activité thermokarstique à l’échelle du bassin versant. L’enregistrement en continu du niveau d’eau et de la turbidité a permis de démontrer que le régime hydrologique et sédimentaire de la rivière Sheldrake est corrélé aux variations des températures de l’air et aux précipitations estivales. Le débit estival est étroitement lié aux chutes de pluie qui favorisent le transport des sédiments (objectif 5). La contribution majeure de ce chapitre est la mise en évidence d’une période de turbidité en 2010 et 2013 survenant en juillet et donc indépendante de la crue printanière mais associée à l’élévation des températures de l’air en début d’été et dans une moindre mesure aux précipitations. Cette période correspond au moment où la couche active progresse jusqu’à ~90% de sa profondeur totale, favorisant l’activité des ostioles et des glissements de terrain (objectif 6). Les sédiments ainsi disponibles sont érodés par le ruissellement et transportés jusqu’au réseau de drainage. Il est donc proposé pour la première fois que le pic de transport de sédiment dans les rivières subarctiques peut être influencé par la température de l’air et le régime thermique de la couche active et de ce fait, indépendamment du pic de débit lié à la fonte des neiges.

De manière générale, le principal défi de tels travaux reste l’accessibilité du terrain, puisqu’on ne peut se rendre sur les lieux qu’en hélicoptère. L’intérêt des capteurs de niveau d’eau et de turbidité fut justement de remédier à cette lacune. Cependant, un plus grand nombre d’échantillons d’eau récoltés ainsi que de mesures de courant opérées renforcerait les conversions et les corrélations. Comme mis en évidence dans le 2ème chapitre, l’état de dégradation du pergélisol du bassin versant de la rivière Sheldrake semble désormais trop avancé pour quantifier les impacts maximaux des rejets de sédiment et de carbone organique dans l’environnement fluvial. Autrement dit, ces impacts ont probablement été plus importants par la passé, comme c’est le cas aujourd’hui dans la région de la rivière

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Nastapoka, située à environ 30 km plus au nord. L’intense activité thermokarstique que subissent certains sous-bassins de la rivière Nastapoka reflète très certainement les conditions qui régnaient il y a quelques décennies sur le bassin de la rivière Sheldrake. Par conséquent, il serait très pertinent d’étendre la méthodologie de recherche à plusieurs cours d’eau le long du gradient climatique latitudinal.

La troisième partie de l’étude (chapitre 4) décrit dans un premier temps la bathymétrie et le relief subaquatique du Passage de Nastapoka au large de l’embouchure de la rivière Sheldrake (objectif 7). On démontre que le relief du fond marin est accidenté avec la présence de deux crêtes asymétriques dans le centre de la zone et des bassins profonds, favorables aux accumulations sédimentaires. L’étude sédimentologique montre la complexité des processus de sédimentation et l’influence des courants de fond, caractérisée notamment par l’absence de sédiment colloïdaux dans la plupart des carottes et la présence de queues de comète et de chenaux d’érosion sur le fond marin. Des structures sédimentaires sont décrites. Une carotte montre des convolutes tronqués par une surface d’érosion interprétés comme étant le résultat de la liquéfaction de sédiments glaciomarins/paraglaciaires pendant la période de transition entre des conditions glaciaires et postglaciaires. Les vitesses de sédimentation sont faibles et variables, proches d’un ordre de magnitude, mais néanmoins en adéquation avec les autres études réalisées en baie d’Hudson. Les taux de sédimentation les plus élevés sont proches des côtes du Québec et de l’Île Gillies et les plus faibles dans les bassins profonds. Les analyses isotopiques et élémentaires du carbone et de l’azote montrent que la fraction de matière organique d’origine terrigène dans les sédiments de fond a augmenté de 30% depuis le Petit Âge Glaciaire et que ce phénomène s’est considérablement accéléré vers la fin du 20ème siècle (objectif 8). S’il est suggéré que la dégradation du pergélisol est un des facteurs qui contribue à cette augmentation, le contexte de changement environnemental (volume des précipitations de pluie et de neige, intensité des débâcles printanières, hydrodynamisme marin, production primaire accrue, …) semble trop large et trop complexe pour que cette augmentation puisse être attribuée uniquement à cette cause.

Étant donné la complexité des reliefs et des courants de fond dans le Passage de Nastapoka, il est probable que les processus de sédimentation et d’érosion varient considérablement sur

131 une courte distance. L’utilisation d’un sonar multifaisceaux et de la sismique-réflexion via un profileur de sous-surface permettrait de préciser avec une haute résolution le cadre particulier de chaque carotte de sédiment, ce qui renforcerait leur interprétation. Par ailleurs, les analyses élémentaires et isotopiques n’ont pas permis de discriminer les sédiments organiques issus de la dégradation du pergélisol en particulier. L’intense hydrodynamisme du Passage de Nastapoka engendre très certainement une large dispersion du matériel fluvial. Pour y remédier, il serait intéressant d’extraire des carottes de sédiment directement sur les lobes sédimentaires de petits cours d’eau de la région, comme la rivière Sheldrake.

En somme, le réchauffement climatique depuis la fin du Petit Âge Glaciaire, accéléré depuis la fin du 20ème siècle, provoque une dégradation importante du pergélisol. À l’intérieur du bassin de la rivière Sheldrake, cette dégradation se fait sentir davantage dans la toundra forestière, à l’est de la limite des arbres. À la désintégration des palses, des lithalses, des plateaux palsiques et des plateaux de pergélisol s’ajoutent les nombreux glissements de terrain et l’érosion des berges fluviales. Avec la dégradation du pergélisol, le réseau fluvial se ramifie un peu davantage, ce qui facilite l’évacuation des sédiments et du carbone via le cours d’eau principal. Découlant du thermokarst, la charge sédimentaire en suspension dans le cours d’eau est plus importante en été alors que les températures élevées de l’air commandent le dégel des sols et créent de l’instabilité sur les versants. Les pluies estivales servent à alimenter les pointes de débit qui évacuent ces sédiments jusqu’à la mer. Une fois rendus à la mer, les sédiments et le carbone transportés en suspension subissent une forte dispersion à cause de l’intensité des courants marins, notamment celui qui s’écoule vers le nord dans le Passage de Nastapoka. Il en résulte l’absence d’une augmentation mesurable du taux de sédimentation; par contre la composition isotopique du carbone sédimentaire suggère fortement un apport accru de la composante terrigène à laquelle la dégradation du pergélisol a dû contribuer, quoique ce ne soit pas le seul facteur qu’on puisse invoquer.

132

Appendix 1: Detailed statistics on permafrost coverage in 1957 and 2009, and changes.

Surficial Lithalsa coverage Lithalsa coverage Change in lithalsa Palsa coverage Palsa coverage Change in palsa Total permafrost Zone Biome deposits in 2009 in 1957 coveragea in 2009 in 1957 coveragea changea m2 m2 %b m2 %b m2 % m2 %b m2 %b m2 % m2 %

Silty clay deposits 2 ST 783,309 489,248 62 493,535 63 -4287 -1 0 0 0 0 0 0 -4287 -1 3 ST 339,325 88,454 26 95,121 28 -6667 -7 0 0 0 0 0 0 -6667 -7 4 ST 1,324,046 534,362 40 557,506 42 -23,144 -4 0 0 0 0 0 0 -23,144 -4 5 ST 996,317 319,689 32 330,397 33 -10,708 -3 0 0 0 0 0 0 -10,708 -3 6 ST 387,029 98,734 26 100,258 26 -1524 -2 0 0 0 0 0 0 -1524 -2 7 ST 592,809 145,288 25 152,192 26 -6904 -5 0 0 0 0 0 0 -6904 -5 8 ST 1,654,949 543,651 33 610,180 37 -66,529 -11 8841 0.5 10,051 0.6 -1210 -12 -67,739 -11 9 ST 1,047,638 192,939 18 200,420 19 -7481 -4 0 0 0 0 0 0 -7481 -4 10 ST 381,057 86,039 23 104,335 27 -18,296 -18 0 0 0 0 0 0 -18,296 -18 11 ST 865,701 121,106 14 147,733 17 -26,627 -18 211 0 369 0 -158 -43 -26,785 -18 12 ST 1,319,777 478,453 36 533,102 40 -54,649 -10 0 0 0 0 0 0 -54,649 -10 14 FT 777,675 114,377 15 155,233 20 -40,856 -26 5864 0.8 14,891 1.9 -9027 -61 -49,883 -29 15 FT 605,272 108,802 18 150,027 25 -41,225 -28 3757 0.6 7189 1.2 -3432 -48 -44,657 -28 16 FT 547,483 36,338 7 41,781 8 -5443 -13 0 0 429 0.1 -429 -100 -5872 -14 17 FT 1,056,487 184,691 18 276,568 26 -91,877 -33 3653 0.3 7218 0.7 -3565 -49 -95,442 -34 18 FT 422,055 14,034 3 14,600 4 -566 -4 0 0 0 0 0 0 -566 -4 19 FT 452,814 1668 0 6601 2 -4933 -75 0 0 0 0 0 0 -4933 -75 20 FT 280,293 36,814 13 46,876 17 -10,062 -22 0 0 0 0 0 0 -10,062 -21 21 FT 550,941 86,711 16 139,136 25 -52,425 -38 297 0.1 865 0.2 -568 -66 -52,993 -38 22 FT 217,524 20,505 9 41,786 19 -21,281 -51 321 0.1 539 0.2 -218 -40 -21,499 -51 23 FT 277,145 5858 2 22,359 8 -16,501 -74 1416 0.5 3710 1.3 -2294 -62 -18,795 -72 24 FT 457,789 90,714 20 134,025 29 -43,311 -32 754 0.2 8735 1.9 -7981 -91 -51,292 -36 25 FT 510,529 59,107 12 83,062 16 -23,955 -29 0 0 4724 0.9 -4724 -100 -28,679 -33 26 FT 643,384 9273 1 39,132 6 -29,859 -76 0 0 3230 0.5 -3230 -100 -33,089 -78 27 FT 809,407 140,355 17 274,096 34 -133,741 -49 5910 0.7 15,343 1.9 -9433 -62 -143,174 -49 28 FT 637,172 19,014 3 48,741 8 -29,727 -61 4692 0.7 12,757 2 -8065 -63 -37,792 -61 29 FT 312,332 2923 1 16,270 5 -13,347 -82 6248 2 19,857 6.4 -13,609 -69 -26,956 -75 30 FT 652,542 20,813 3 76,066 12 -55,253 -73 1005 0.2 5795 0.9 -4790 -83 -60,043 -73 31 FT 79,084 0 0 783 1 -783 -100 0 0 0 0 0 0 -783 -100 32 FT 525,762 50,954 10 116,788 22 -65,834 -56 0 0 4960 0.9 -4960 -100 -70,794 -58 33 FT 448,957 7683 2 23,575 5 -15,892 -67 1638 0.4 7705 1.7 -6067 -79 -21,959 -70 34 FT 579,853 57,223 10 101,375 18 -44,152 -44 3393 0.6 11,063 1.9 -7670 -69 -51,822 -46 Total 14,458,669 1,946,394 14 2,794,470 19 -848,076 -30 39,159 0.3 129,379 0.9 -90,220 -70 -938,296 -32,0 Peat deposits 37 FT 161,168 6578 4 12 7 -5360 -45 18,989 11.8 30,949 19.2 -11,960 -39 -17,320 -40 38 FT 141,641 0 0 1747 1 -1747 -100 17,015 12 27,845 19.7 -10,830 -39 -12,577 -43 39 FT 145,575 7223 5 24 16 -16,649 -70 12,450 8.6 19,143 13.1 -6693 -35 -23,342 -54 41 FT 227,978 1265 1 14 6 -12,365 -91 16,471 7.2 36,061 15.8 -19,590 -54 -31,955 -64 Total 676,362 15,066 2 51,187 8 -36,121 -71 64,925 9.6 113,998 16.9 -49,073 -43 -85,194 -52

TOTAL 21,212,815 4,180,886 19.7 5,184,846 24.4 -1,003,960 -19.4 112,925 0.5 253,428 1.2 -140,503 -55 -1,144,463 -21

a Between 1957 and 2009 b Percentage of permafrost coverage in the area covered by surficial deposits

133

Appendix 2: Detailed statistics on thermokarst ponds coverage in 1957 and 2009, and

changes.

TPa change CTPb change CTPPc in Zone TPa in 1957 TPa in 2009 1957-2009 CTPb in 1957 CTPb in 2009 1957-2009 2009 m2 %d m2 %d m2 % m2 %d m2 %d m2 % m2

2 1384 0.2 1407 0.2 23 2 0 0 0 0 0 0 0 3 1579 0.5 1680 0.5 101 6 0 0 0 0 0 0 0 4 8932 0.7 11,267 0.9 2335 26 0 0 0 0 0 0 0 5 7425 0.7 8276 0.8 851 11 2701 36 2866 35 165 6 0 6 376 0 2270 0.6 1894 504 168 45 259 11 91 54 0 7 70 0 464 0.1 394 563 0 0 0 0 0 0 0 8 20,482 1.2 22,538 1.4 2056 10 0 0 0 0 0 0 0 9 46,985 4.5 93,676 8.9 46,691 99 13,791 29 26,358 28 12,567 91 913 10 2065 0.5 5643 1.5 3578 173 156 8 457 8 301 193 0 11 10,604 1.2 20,944 2.4 10,340 98 4865 46 6971 33 2106 43 0 12 21,555 1.6 42,464 3.2 20,909 97 3983 18 7526 18 3543 89 0 14 25,922 3.3 65,773 8.5 39,851 154 5029 19 11,432 17 6403 127 2821 15 21,658 3.6 50,022 8.3 28,364 131 7173 33 10,449 21 3276 46 2698 16 3033 0.6 5982 1.1 2949 97 514 17 1124 19 610 119 0 17 68,561 6.5 139,768 13.2 71,207 104 34,856 51 56,973 41 22,117 63 0 18 0 0 338 0.1 338 \ 0 0 217 64 217 0 0 19 1793 0.4 3714 0.8 1921 107 371 21 994 27 623 168 0 20 2253 0.8 4072 1.5 1819 81 809 36 1648 40 839 104 0 21 10,022 1.8 25,279 4.6 15,257 152 920 9 1499 6 579 63 127 22 9024 4.1 16,848 7.7 7824 87 2734 30 4261 25 1527 56 123 23 12,090 4.4 20,210 7.3 8120 67 3938 33 5397 27 1459 37 1010 24 36,719 8 67,746 14.8 31,027 84 13,638 37 15,357 23 1719 13 1470 25 30,778 6 63,494 12.4 32,716 106 9376 30 17,804 28 8428 90 1442 26 12,147 1.9 24,052 3.7 11,905 98 2766 23 4599 19 1833 66 0 27 56,257 7 124,517 15.4 68,260 121 7391 13 11,141 9 3750 51 0 28 11,159 1.8 19,914 3.1 8755 78 2180 20 2984 15 804 37 1110 29 10,121 3.2 15,109 4.8 4988 49 5790 57 7076 47 1286 22 1913 30 28,778 4.4 51,987 8 23,209 81 2869 10 3599 7 730 25 0 31 730 0.9 949 1.2 219 30 0 0 0 0 0 0 0 32 35,843 6.8 70,810 13.5 34,967 98 8684 24 9798 14 1114 13 860 33 6413 1.4 13,868 3.1 7455 116 1154 18 2687 19 1533 133 0 34 13,210 2.3 32,724 5.6 19,514 148 6435 49 11,222 34 4787 74 1540 37 10,896 6.8 16,597 10.3 5701 52 5753 53 8440 51 2687 47 3386 38 5862 4.1 8358 5.9 2496 43 3068 52 3866 46 798 26 1140 39 11,981 8.2 20,732 14.2 8751 73 0 0 0 0 0 0 0 41 10,069 4.4 19,849 8.7 9780 97 1521 15 4500 23 2979 196 957 Total 556,776 2.6 1,093,341 5.2 536,565 96 152,633 27 241,504 22 88,871 58 21,510

a Thermokarst pond coverage b Connected thermokarst pond coverage c Area covered by connected thermokarst ponds coming from palsa collapse

134

d Percentage of thermokarst pond coverage in the area covered by surficial deposits

135