Documenting the Response of Newfoundland Forests to The

Home , Preboreal

The Marine Record of Abrupt Climate Change at Bay of Islands, Newfoundland

Simone Sandercombe Department of Geography McGill University Montreal, Quebec December, 2011

A thesis submitted to McGill University in partial fulfillment of the requirement of the degree of Masters of Science

Table of Contents

Abstract vi

Résumé vii

Acknowledgements viii

List of Figures iv

List of Tables v

Chapter 1. Background and Literature review 1

Introduction 1

Chapter 2. Study area and Methodology 13

Study area 13

Methods 15

Core description 15

Laboratory procedure 17

Analysis 17

Chapter 3. Results 20

Palynomorph Concentrations 20

Comparaisons with other pollen and dinoflagellate cyst records in the area 20

Pollen assemblage zones 24

Dinoflagellate cyst assemblage zones 26

Reconstructions of sea surface conditions 27

Reconstructed air temperature 29

Chapter 4. Discussion 31

Younger Dryas 31 ii

Preboreal Oscillation 35

8.2 ka event 37

Chapter 5. Summary and Conclusions 40

References 43

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List of Tables

Table 1.Current tree species present around Bay of Islands 51

Table 2. Radiocarbon dates for core MD99-2225, Bay of Islands 52

Table 3. Summary of radiocarbon dates 53

Table 4.Summary of Pollen assemblage data for Bay of Islands, Joes Pond, Southwest Brook Lake, Robinsons Pond, Northern Baie Verte Peninsula and Compass Pond 56

Table 5.Summary of dinoflagellate cyst data for Bay of Islands, St. Anne’s Basin and Southern Labrador Sea. 57

List of Figures

Figure 1. Location map of Bay of Islands, Newfoundland. 58

Figure 2.Core stratigraphy for core MD99-2225, Bay of Islands 59

Figure 3: Concentrations of palynomorphs per gram of sediment in core MD99-2225, Bay of Islands 60 Figure 4.Pollen diagram for core MD99-2225, Bay of Islands 61

Figure 5.Dinoflagellate cyst diagram for core MD99-2225, Bay of Islands 62

Figure 6.Reconstructions of sea surface conditions for core MD99-2225, Bay of Islands 63

Figure 7. Reconstructed air temperatures for core MD99-2225, Bay of Islands 64

Abstract

This study assesses vegetational response to abrupt climatic changes that occurred between ca. 12,000 and 8,000 yr BP using a palynological record from Bay of Islands (west coast of Newfoundland). Western Newfoundland is located near the boundary of two major ecozones: the boreal forest and the tundra. This transition zone is very sensitive to climate change. Core MD99-2225 was extracted from Humber Arm at a depth of 104m, roughly 12km from the Humber River. Samples were analysed every 10cm between 10 and 25 m down-core. Pollen analysis was used to track the evolution of vegetation, and the proportions of both tundra and more thermophilous hardwood taxa were used to determine changes in ecotone. The pollen record was also used to reconstruct air temperatures. The dinoflagellate cyst record was used to reconstruct sea surface parameters, such as temperature, salinity and sea ice.

Results reflect large shifts in the composition of the vegetation and of dinoflagellate cysts assemblages in response to three cold climatic events. At the onset of the Younger Dryas, air temperatures dropped by 5°C from-15°C in February and 16 °C in August , August SST dropped by 10°C from roughly 15°C, and the duration of the sea ice cover increased to 10 months yr-1. Conditions remained harsh between ca. 10,800-10,300 yr BP, Vegetation was dominated by shrubs and grasses and dinoflagellate cyst assemblages are characterized by low species diversity and dominance by arctic species, such as Brigantedinium spp. Sea surface and air temperatures improved following after the Younger Dryas.

At ca. 9,000 yr BP, sea surface conditions and air temperatures decreased in the bay as a response to the Preboreal Oscillation. Shrubs and grasses re-invaded the region and dinoflagellate cyst assemblages showed a decrease in diversity and dominance by arctic species such as Brigantedinium spp.

The last event to be detected in this record, the 8.2 ka cold period, had a smaller impact in the region. Sea surface temperatures as well as air temperature decreased, but for a short time. The proportion of shrub pollen increased slightly and cold water species (Brigantedinium spp., I. minutum and P. dalei) dominated the dinoflagellate cyst assemblages.

Keywords : abrupt climate change, dinoflagellate cysts, vegetation change, sea surface temperatures

Resume

Cette étude évalue la réponse de la végétation aux changements climatiques abrupts survenus entre 12,000 and 8000 ans BP, en utilisant un enregistrement palynologique de Bay of Islands sur la côte ouest de Terre Neuve. Cette région est à la frontière entre deux écozones majeures: la forêt boréale et la toundra. Cette zone de transition est très sensible aux changements climatiques. La carotte MD99-2225 a été prélevée à une profondeur de 104 m à environ 12 km de la rivière Humber. Des échantillons ont été analysés à intervalles de 10cm entre 10 et 25 m de profondeur dans la carotte. L’analyse pollinique a permis de retracer l’évolution de la végétation, et les proportions de taxons caractéristiques de la toundra et celles des arbres plus thermophiles ont servis à déterminer les changements d’écotones. Les données polliniques ont aussi permis de reconstituer les températures. L’enregistrement des kystes de dinoflagellés a permis de reconstituer les paramètres des eaux de surface, tels que la température, la salinité et la durée du couvert de glace saisonnier.

Les résultats montrent d’importants changements dans la composition de la végétation et dans les assemblages de kystes de dinoflagellés, en réponse à trois épisodes de refroidissement climatiques. Au début du Dryas récent, les températures de l`air qui étaient de -15°C en février et de 16°C en août ont chuté de 5°C. Les températures des eaux de surface qui étaient de 15°C ont également chuté de 10°C, et la durée saisonnière du couvert de glace marin a augmenté pour atteindre 10 mois/an. Les conditions sont demeurées très froides entre environ 10800-10300 yr BP. La végétation était dominée par des arbustes et des plantes herbacées. Les assemblages de kystes de dinoflagellés étaient caractérisés par une faible diversité spécifique et une dominance par les taxons arctiques. Les températures de l’air et des eaux de surface se sont améliorées après le Dryas récent. Autour de 9000 yr BP, une autre détérioration climatique a été enregistrée en réponse à l’oscillation PréBoréale. Les arbustes et les plantes herbacées ont de nouveau envahi la région et les assemblages de kystes de dinoflagellés ont connu une baisse de diversité et une augmentation de la proportion de taxons arctiques.

Le dernier événement climatique détecté dans cet enregistrement, le refroidissement de 8.2 ka, a eu un impact moindre. Les températures de l’air et des eaux de surfaces ont connues une baisse, mais pour un lapse de temps très court. La proportion de pollen d’arbustes a augmenté légèrement, de même que celles des espèces indicatrices de conditions d’eaux de surface froides (Brigantedinium spp., I. minutum and P. dalei).

Mots clés: changement climatiques abrupts, kystes de dinoflagellés, changement de végétation, températures des eaux de surface.

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Acknowledgements

I am very appreciative of all the support and guidance I received while writing this thesis. First of all I would like to offer my deepest gratitude to my co supervisors Gail Chmura and Elisabeth Levac; this thesis would not have been possible without their continued encouragement, expertise and advice. To Elisabeth Levac whom I have worked with for 6 years, thank you for all the phone calls, emails, meetings and friendly support, you have been a great friend and mentor.

Thank you to the lab assistants that helped process and organize all of my data, thank you to Eric Fortier, Vanessa Asselin and Thomas Neulieb. I would also like the thank the ship crew on the Marion Dufresne for extracting core MD99-2225 as well as Ali Aksu, from Memorial University, for allowing E. Levac to sample the core. I am very grateful to Matthew Peros for his ongoing help and assistance with C2 software and pollen transfer functions.

A special thanks to my fellow McGill geographers, without you guys I would not have been able to make it through those stressful sleepless nights, especially to Amanda Alfonso, Camille Ouellet Dallaire and Florin Pendea for their friendship, continued reassurance and advice.

Finally I owe my deepest gratitude to my family, without their continued support and encouragement to reach my goals.

I would also like to acknowledge the support of the Natural Science and Engineering Research Council of Canada (NSERC) as well as the support from the Global and Environmental Change Center (GEC3) for funding this research.

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Chapter 1. Background and Literature Review

The climate in Eastern North America over the past 14,000 years was variable, with several periods of near glacial conditions interrupting the post glacial warming trend following the retreat of the Laurentide Ice Sheet: the Younger Dryas, the Preboreal

Oscillation (PBO) and the 8.2 ka event.Northeastern North America is an excellent location to study abrupt climatic changes because it is close to North Atlantic Ocean sites of meridional overturning, and the region`s climate should reflect changes in meridional overturning. Because vegetation is controlled by climate, it should reflect significant changes. My research examines the response of vegetation in western Newfoundland to these abrupt climatic changes. Understanding the response of vegetation to abrupt climate change will help us plan and possibly mitigate potential future impacts of climate change.

To assess vegetation and sea surface condition responses to climate change, interpretations from dinoflagellate cysts and evolution shifts in vegetation in western

Newfoundland were examined during the end of the deglaciation period, between ca.

12,000 and 8,000 years ago, using a high resolution sediment core (MD99-2225) from

Bay of Islands.

Vegetation response to abrupt climate change

Understanding what constitutes an abrupt climate change and how such events influence vegetation is essential to anticipate how vegetation may be affected by similar abrupt changes in the future. But how do we define abrupt climate change? There are many different interpretations of the meaning. The U.S. Climate Change Science Program Subcommittee on Global Change Research (2008) proposed this definition: “a large-scale change in the climate system that takes place over a few decades or less, persists (or is anticipated to persist) for at least a few decades, and causes substantial disruptions in human and natural systems.” The NRC (Natural Resource Council, 2002), published the report “Abrupt Climate Change” in which they offered a mechanistic and an impact-based definitions of abrupt climate change. The mechanistic definition defines abrupt climate change as “occurring when the climate system is forced to cross some threshold, triggering a transition to a new state at a rate determined by the climate system itself and faster than the cause.” The impacts-based definition of abrupt climate change is

“one that takes place so rapidly and unexpectedly that human or natural systems have difficulty adapting to it.” Alley and Agustsdottir (2005) also provide a definition, “Abrupt climate change typically occurs when the climate system is forced across some threshold, causing evolution to a new, more-or-less persistent state at a rate determined by the system and faster than the cause” (p 1124). Although there are different interpretations of abrupt climate change, they all imply that the rates of change are too rapid for natural systems to adapt.

When using vegetation and fossil pollen to reconstruct past climates certain assumptions must be accepted; vegetation responds to changes in climate and because of this, species distributions and abundances are in equilibrium with each other (Davis and

Botkin, 1985). It is also assumed that there are no external factors influencing how each species responds to climate and that the pollen assemblage accurately resembles the vegetation composition (Davis and Botkin, 1985). Davis and Botkin (1985) performed four sensitivity tests in order to understand the factors that alter vegetation distribution 2

and abundance. Using a computer model they tested vegetation’s sensitivity to the length of a climatic event and to different temperatures. Their results showed that the amplitude of temperature change had a larger impact on the vegetation abundance and composition than the duration of the climatic event.

To determine whether the past abrupt changes in temperature were too rapid for postglacial vegetation to adapt Williams et al. (2002) looked at 11 high-resolution postglacial lake records from North America and Europe. Williams et al. (2002) found that vegetation adapted to the rapid climate change events, both warming and cooling, and that the shortest lag time at each site was less than 100 years. There are several explanations for the short lag times in the vegetation response they observed. One is that low tree densities around the tree line may have allowed for more light penetration enabling the saplings to grow and thrive at an increased rate when climate warmed up

(Williams et al., 2002). Another theory is that the mainly herbaceous late glacial plant communities were able to respond more rapidly to climate change than vegetation such as trees which are slower to mature and reproduce (Williams et al., 2002).

Each plant species has a different response to climate change and those individualistic responses will cause changes in plant associations. This explains why the plant associations that are characteristic of today's biomes are not permanent features and why the associations tend to have a shorter life span than the distribution and expansion of individual species (Williams et al., 2004).

There are many advantages to marine records however, interpreting marine pollen can be challenging. The pollen can be transported by the wind, by streams and rivers,

ocean currents, and glacial meltwater streams. Transportation modes are not the only factor determining the composition of the assemblage: the morphology, productivity and release mechanisms of different plant species also affect the percentages and diversity of the assemblage (Mudie and McCarthy, 1994).

Fluvially transported pollen is subject to many obstacles, the grains can be ingested by organisms, deposited and possibly re-suspended on the sides of the river, or carried in ocean currents once the river enters the ocean. Rivers carry pollen from throughout their drainage basin that are derived from both local vegetation as well as long distance transport. Long distance transport of pine pollen is recognized ( Anderson and

Lewis, 1992; McCarthy et al., 2003) and it is likely to be abundant in early parts of the pollen record when local vegetation produced less pollen.

The transport of grains by wind involves many factors such as wind-speed through the trunk space, or through the forest canopy, density of the cover, thickness of the foliage, time of the dispersal, and the size, shape and proximity to the first major water body. All of these will ultimately determine to the composition of the final assemblage (Lowe and Walker, 1997; Davis, 2005). Lowe and Walker (1997) note the constant ‘background’ component that can lead to complications, especially in areas with low plant production, such as the Bay of Islands during cooler periods. The background component is grains that are constantly held within the upper air currents throughout the world that ultimately get deposited very far away from their source.

The last ice age in Newfoundland and the history of deglaciation

Most of Newfoundland was under ice during the last glacial maximum but there is no agreement about the extent of the ice sheet along the Eastern Canadian Margin. In an early study Flint (1940) proposed the ‘maximum’ model of the Late Wisconsin ice cover in which the Laurentide ice sheet was combined with local ice caps over Nova Scotia,

Prince Edward Island and Newfoundland and extended to the edge of the continental shelf. Although Flint's early ‘maximum’ model was questioned by others it was supported by the work of Grant (1969a; 1972). Grant’s work then lead to the development of the ‘minimum’ model of Late Wisconsin ice cover. The ‘minimum’ model describes the Laurentide ice sheet extent to be limited to the mainland of eastern

Canada. The ‘minimum’ model presents the Late Wisconsin ice in the form of several ice caps, with nunataks, rather than a single ice sheet on Newfoundland (Rogerson 1981).

Maps by Grant (1997) and Brooks (1977) provide evidence of at least four successive ice advances in Newfoundland. The maps also show evidence of an independent

Newfoundland ice cap with a few separate centers instead of one mass of ice, consistent with an earlier proposition by Rogerson (1981). Grant's (1987;1989a;1989b;1994) work also strongly supported the Nunatak hypothesis earlier proposed by Fernald (1911) and

Coleman (1926). It implies that although there was ice present on Newfoundland during the LGM, it was not present on the summits of the mountains, thus providing refugia for flora and fauna. This hypothesis has been tested using cosmogenic nuclide exposure dating on the Long Range Mountains, on the west coast of Newfoundland (Osborne et al.,

2007). The results from the tests confirm that a thick warm based Laurentide ice sheet did not cover the island; instead there was a thinner independent ice sheet over 5

Newfoundland, which did not cover the mountain summits, resulting in the existence of nunataks (Osborne et al., 2007). By dating shells from the Bay of Islands region, Grant

(1987) found that the glacier had retreated from the outer coast before 15937±1075 cal yr

BP and was inland of Corner Brook by142616±857 cal yr BP. Grant (1987) also found evidence for the return of glacial like conditions during the Younger Dryas and the PBO.

Ocean patterns and currents are controlled by the strong wind systems over the

North Atlantic Ocean. These wind systems have not remained consistent over time due to the presence and magnitude of glacial ice and therefore the ocean patterns also have been altered. When the deglaciation first started ca. 17,000-15,000 yr BP, orbital changes led to more rapid warming and ice sheet decay (Mudie and McCarthy, 1994). Between ca. 10,000-7,000 yr BP the ice sheet continued to decrease in size and as a result led to the establishment of the modern distributed prevailing westerly surface winds (Ruddiman and Duplessy, 1985).

Evolution of postglacial climate

The Younger Dryas was a period of glacial cooling amidst postglacial warming that started approximately 11,000 yr BP and lasted until 10,000 yr BP (Berger, 1990,

Mayle et al. 1993). The Younger Dryas is, according to Peteet (2000), the best- documented example of an abrupt climate change: within its millennial duration the cooling was so significant that it caused a reversal in the vegetation succession and the re- expansion of the glaciers. In Western Newfoundland this meant the return of tundra vegetation (Anderson and Lewis, 1992; McCarthy et al., 1995). 6

The cause of the Younger Dryas has been heavily debated. Although it was mostly agreed upon that the freshwater discharge from Lake Agassiz led to the cooling, the lack of physical evidence of the flood led to debates. However, in a recent study

Carlson et al., (2007) used geochemistry of planktonic foraminifera collected from the mouth of the St. Lawrence estuary to evaluate whether freshwater was released during the

Younger Dryas and whether the influx was large enough to cause a slowing of thermohaline circulation. Carlson et al., (2007) found that there was in fact a large enough influx, 0.06 ±0.02 Sverdrup, at the start of the Younger Drays to have affected the thermohaline circulation.

There is now consensus that the drainage of Lake Agassiz and the large influx of freshwater were responsible for weakening the formation of North Atlantic Deep water

(NADW) resulting in cooler temperatures around the North Atlantic (Anderson and

Lewis, 1992; Alley et al., 1997; Barber et al., 1999; Drinkwater et al., 1999; Broeker

2000; Bard, 2001; Maslin et al., 2004; Mayewski et al., 2004; Schulze et al., 2004; Alley and Agustsdottir, 2005; Wiersma and Renssen, 2006; Anderson et al., 2007; Carlson et al., 2007; Kleiven et al., 2008; Thomas et al., 2007). This cooling period is well document within sediment cores around the North Atlantic as well as in ice cores from

Greenland. In Western Newfoundland, the Younger Dryas is recorded as a vegetation reversal (Anderson and Macpherson 1994) and is responsible for the re-advance of ice caps (Grant 1989). Evidence of the event is throughout Eastern Canada (Mott, 1986;

Mayle and Cwynar 1995; Levesque et al., 1997; Yu and Eicher, 1998; Miller and Elias,

2000; Broecker, 2006), Yu and Eicher (1998), but also at two sites in Ontario, Twiss Marl

Pond and Crawford Lake. At both sites shrub and herb pollen increased relative to tree 7

pollen, suggesting more open forests thus cooler temperatures. Yu and Eicher (1998) noted the Younger Dryas at their two sites by a decrease in δ18O of lake carbonates.

Brand and McCarthy (2005) also measured decreased δ18O values in mollusk shells indicating a cooling period, from two other sites in Ontario, Navan and Bearbrook.

The PBO was a cold event that lasted about 150-250 yrs (Fisher et al., 2002).

Although Greenland ice core records two cold events (between 10,000-9,900 and 9,600 -

9,500 yr BP), Bjorck et al. (1996, 1997) suggest that the two fluctuations represent one event, the PBO. The PBO is also recorded in some high-resolution climate records within the North Atlantic region; however, because of its short duration, it is not detectable within all records. Anderson et al. (2007) note the inconsistency of the PBO within pollen records in eastern North America and suggest that records showing vegetation responses were located close to an ecotone, while those records showing very little to no changes were not. According to Peteet (2000), “vegetational response is best expressed in terrestrial records near ecotones, where sensitivity to climate change is greatest, and response times are as short as decades.” This emphasizes the importance of site location, especially when looking for responses of abrupt short-lived events, such as the PBO. As with the Younger Dryas, the PBO was detected at other sites with high resolution records, such as Crawford Lake and Twiss Marl Pond whereYu and Eicher (1998) found a slight decrease in the δ18O record at ~ 9600 yr BP. They also noted an increase in Picea pollen at both sites during this period, associating both isotope and pollen shifts with cooling.

The 8.2 ka event was another brief cold event that lasted between 200-400 years

(Alley et al. 1997; Alley and Agustsdottir 2005;Thomas et al., 2007). It was first found

within the Greenland ice core records as reduced δ18O (Thomas et al., 2007). There is a consensus that the cooling occurred due to the final drainage of Glacial Lake Agassiz, and that the resulting influx of freshwater into the North Atlantic weakened the thermohaline circulation leading to cooling (Barber et al., 1999; Alley et al., 1997; Alley and Agustsdottir, 2005; Thomas et al., 2007). Lewis et al. (2009) recently revised the age of the event to 8.33 ka with an error range of 8.15-8.48 ka. This period of cooling, although short-lived, was very significant. According to Alley et al. (2007), the 8.2 ka event was the most prominent Holocene event with about half the amplitude of the

Younger Dryas. Analyses of sediment core data, ice core data and experiments with climate models indicate that temperatures dropped between 1.5-3.0°C around the northeastern North Atlantic Ocean and that the cooling caused major displacement of vegetation patterns (Alley et al., 1997; Barber et al., 1999; Alley and Agustsdottir, 2005;

Thomas et al., 2007; Levac et al., 2011). Although the 8.2 ka event clearly stands out within the Greenland ice cores, due to its short duration, it is not always represented within sediment records from North America unless they have high temporal resolution

(Alley and Agustsdottir, 2005). Indeed, in the resolution cores from Crawford Lake Yu and Eicher (1998) found a decrease of about 0.8‰ in the δ18O record, at ~8200 yr BP which suggested a cooling of about half the magnitude of the Younger Dryas. Shuman et al. (2002) recorded the event in Pennsylvania with increased percentages of Picea and

Betula. Seppa et al. (2002) also found pollen responses in Nunavut with decreased percentages in Picea and an expansion of the tundra into tree covered regions for approximately 200 years.

Postglacial migrations of vegetation

Once the ice retreated vegetation became established and primary succession began. The physical landscape was continually altered throughout the Holocene due to the abrupt changes in climate, some caused by large discharges of meltwater to the North

Atlantic, leading to glacier re-advance (as seen above). The two biomes located near Bay of Islands, the boreal forest and the tundra, have undergone significant changes in composition, abundance, and distribution through time in response to climate change and ice retreat. The largest changes occurred in the early Holocene between 11,000 and 7,000

BP (Bernabo and Webb, 1977), as the Laurentide ice sheet was receding rapidly and more land was being uncovered allowing for more plant growth.

At ~14,000 yr BP ice recession was minimal; the ice was still close to its maximum extent (Dyke, 2005). As the ice receded between ca. 13,500 and 11,000 yr BP the land just south of the Newfoundland ice sheet became vegetated by a shrub- sedge- tundra landscape, with Salix, Betula, Juniperus, and Artemisia,along with sedges, grasses and ericaceous plants (Anderson and Macpherson, 1994). The shrub tundra in western and southwestern Newfoundland implies either relatively rapid warming after glaciation or rapid plant migration. The warming was significant enough for a more diverse vegetation cover to become established, but the more temperate species did not have sufficient time to migrate northward (Anderson and Macpherson 1994).

There were several ice re-advances during the postglacial period that lead to the development of major end moraine systems. Both the Piedmont Moraines and

Robsinson’s Head Drift can be found along the outer west coast of Newfoundland (Dyke

2005). For the most part however, along the west coast of Newfoundland, the ice was retreating inland ca. 12,000 yr BP. On the eastern coast, warming did not progress as fast as on the western coast, the land was sparsely vegetated with herbs and the ice remained in contact with the marine water until 12,000 yr BP (Anderson and Macpherson, 1994).

Between ca. 11,000 and 10,000 yr BP the shrub-tundra on the western coast of

Newfoundland was replaced by a herb tundra, which represents a reversal in the expected vegetation succession. This period coincides with the Younger Dryas cooling. Although the ice had started to retreat, the abrupt climatic cooling lead to re-advancement of the ice

(Anderson and Macpherson, 1994). The cooling was so significant on Newfoundland that the ice on the Long Range Mountains of the Northern Peninsula re-advanced into the

Goldthwait Sea and deposited the Ten Mile Lake Moraine, which was dated post ca.

11,000 yr BP (Grant, 1989). Cooling also affected the eastern and western coasts of the island with the formation of ice-wedges that are believed to be younger than ca. 11,000 yr

BP (Anderson and Macpherson, 1994).

The ice re-advance however, did not last long and there was once again a northern migration of shrub tundra and several boreal tree species that migrated into the area. After

~10,000 yr BP the interior of Newfoundland was ice free and the relative sea level rose

(Shaw, 2005). Picea became established on the island by ~10,000 yr BP and in western

Newfoundland by ~9500 yr BP (Anderson and Lewis, 1992; Macpherson, 1995). During the PBO, the forests were once again replaced by shrub tundra vegetation on the west coast of the island (Macpherson, 1995). Although this cold period was significant enough to cause a reversal in the vegetation succession, it did not persist for long, about 200

years (Cwynar et al., 2003), and temperatures once again began to increase allowing for the migration of Pinus on the Island. Pinus became established in western Newfoundland by ca. 8,500 yr BP (Macpherson, 1995). As the temperatures continued to increase the modern day boreal forest vegetation started to become established on the west coast of the island and by ca. 7,000 yr BP the forests as we know them today were established.

Since then, Picea has been the most prominent species in the region followed closely by

Abies balsamea and Betula (Dyke, 2005).

Chapter 2. Study Area and Methodology

Study Area

The Bay of Islands is located on the west coast of Newfoundland (Figure 1) and it has an expanse of roughly 200 km2, with water depths reaching 200 m. The regional topography around the bay rises 800 m above sea level. Core MD99-2225 was extracted in 1999 by the French research vessel “Marion Dufresne” from Humber Arm, located in the east side of the bay at a water depth of 104 m, about 12 km from the Humber River

(48°59.88’N, 58°05.08'W). The Humber River is approximately 120 km long and flows through Deer Lake before entering the bay. This location was chosen because of its close proximity to the mouth of the Humber River and the resulting high sedimentation rate which enables higher temporal resolution compared to many other marine records.

Bay of Islands is a sub-basin of the highly stratified Gulf of St. Lawrence. The

Gulf has a major fresh water influx from the St. Lawrence River, approximately 10,100 m3s-1 (Smith and Conover, 2002). The water exchanges between the Bay and the Gulf are an important factor in sea surface temperature and salinity within the Bay of Islands (de

Vernal et al., 1993; Levac, 2003).

Although Humber arm is located within the Bay of Islands it is still largely influenced by the Labrador Current. The Labrador Current is a southerly flowing cold ocean current in the North Atlantic Ocean which originates in Davis Strait and the

Labrador Sea and flows along the coast of Labrador, around Newfoundland and along the east coast of Nova Scotia (Mudie and McCarthy, 1994). Not only is this area of

Newfoundland influenced by the Labrador Current, but also the Gulf of St. Lawrence and 13

Cabot Strait, located to the west and Southwest, also significantly affect the sea surface conditions as well as other climatic properties (Banfield, 1983). In the winter months the sea surface temperature (SST) is below 0°C throughout the Gulf of St. Lawrence while in the summer the temperature is more variable, ranging from 12 to 16°C (Banfield, 1983).

Because Bay of Islands is highly influenced by the Gulf the sea surface temperatures are very similar, during the winter months the average temperature is -0.4°C and during the summer the temperature can reach up to 16.7°C (NOAA, 1994). Present day sea surface salinity (SSS) in the Gulf of St. Lawrence is 34.1 ± 0.7 practical salinity units (psu (de

Vernal and Hillaire-Marcel, 2000). During the winter months the west coast of

Newfoundland experiences periods of sea ice cover, in Bay of islands sea ice is present for two months per year, on average (Levac, 2003). According to Drinkwater et al.

(1999), the ice first arrives near the end of January and is present through to the end of

March.

The land around the Bay of Islands, due to the protection from the Long Range

Mountains, experiences a warmer climate than the rest of the Island. The western

Newfoundland ecoregion is significantly drier than other regions on the island due to the rain shadow effect from the Long Range Mountains (Banfield, 1983). Corner Brook receives an average of 99.3 mm of precipitation in February and 98.6 mm in August

(Environment Canada, 2011). Because of the mountains' protection, Humber arm experiences frequent calm winds and the rare fog event, and it has the highest daily mean temperature on the west coast of the island (Banfield, 1983). The average air temperature at Corner Brook in February is -7.2°C and 16.9°C in August (Environment Canada,

2011). 14

These climate characteristics make the western Newfoundland eco-region which surrounds the Bay of Islands, the western Newfoundland eco-region, unique to the Island because it has the most favorable conditions for plant growth such as, fertile soils, higher mean annual temperature and daily maxima as well as protection from the cold north and east winds by the mountains (Damman, 1983).

The western Newfoundland ecoregion is heavily dominated by Abies balsamea, however, other coniferous and deciduous tree species are also present (Table 1). Betula lutea, Pinus strobus, Picea mariana and Picea glauca are present in substantial amounts together with small stands of Populus tremuloides and shrub Betula (Damman, 1983).

Within the western Newfoundland eco-region the following tree species reach their northern limit of distribution: Pinus strobus, Acer rubrum, and Populus tremuloides. This is also the only ecoregion on Newfoundland in which Fraxinus nigra is present, and where Acer spicatum is resilient enough to compete with other species (Damman, 1983).

Methods

Core description

Core MD99-2225 is 37.53 m long, however only the depths between 1000– 2500 cm were analysed for this study (Figure 2), these depths were chosen because, through comparison and calculation, they represent the three targeted abrupt climatic events. The upper part of the core (0-1475 cm) is composed of dark grey to black clayey silt, and bioturbation is locally important. The section between 1470-1660 cm is composed of

banded clayey silt in colours of grey and reddish brown; with local laminations and graded sequences. The section between 1660- 2540 cm is composed of bioturbated clayey silt similar to the upper section (Hillaire-Marcel et al., 1999). The section between 1450-

1550 cm shows banded silty clay sections with reddish brown bands.

Two levels (1440 cm and 2166 cm) within core MD99-2225 (Table 2) were dated by Levac (2003) who used hinged mollusc shells. These dates were recalibrated with the program CALIB 6.0.1 (Reimer et al., 2009; Stuiver and Reimer, 1993) using the

MARINE 09 calibration dataset (Reimer et al., 2009) with a delta R value of 150 ± 42.

Dates from other studies in Eastern Canada were used to constrain ages in the MD99-

2225 core, but re-calibrated (Table 3.). The terrestrial dates on lake organic sediment were calibrated with the CALIB 6.0.1 program (Reimer et al., 2009; Stuiver and Reimer,

1993) using the INTCAL09 dataset (Reimer et al., 2009) while the two marine cores were calibrated using the MARINE 09 dataset (Reimer et al., 2009) with calculated delta R values of 121 ±17 for La Have Basin and 83 ±14 for St. Anne’s Basin (Table 2). Ages between radiocarbon dates were linearly interpolated.

From these interpolations the sedimentation rate is estimated to be highest in the middle of the core. Between 3000 yr BP and 9630 yr BP the sedimentation rate is 1.4 mm yr-1, between 9630 cal yr BP and 9990 cal yr BP the sedimentation rate is 3.86 mm yr-1, between 9990 yr BP and 11,910 yr BP the sedimentation rate is 1.87 mm yr-1 and between 11,910 yr BP and 13,290 yr BP the sedimentation rate is 1.08 mm yr-1.

According to these calculations, the age at the base of the studied section of the core

(2500cm) is ~16000 yr BP and the top of the studied section of the core (1000cm) is

~7500 yr BP.

Laboratory procedures

A total of 96 samples were taken every 10 cm between 10 and 25 m down core.

Each sample was 5 cm3. Exceptions were the samples taken at 1040-1042 cm and 1060-

1062 cm which were 4 cm3 and the samples taken at 1070-1072 cm and 1080-1082 cm which were 4.5 cm3. Samples were sieved through a 120 and 10 μm sieve and the materials retained on the 10 μm sieve was used for analysis. After sieving one

Lycopodium spore tablet (18,584 spores/tablet), from batch 177745 was added to each sample after sieving to allow calculation of palynomorph concentrations(Maher, 1981;

Stockmarr, 1971), pollen accumulation rates (PAR) and pollen-influx (PI). The concentrations were calculated as the number of palynomorphs counted times total spores in a tablet/ number of spores counted / sample volume.

Samples were treated with hot 49% hydrochloric acid and hot 10% hydrochloric acid to remove carbonates and silicates, respectively (Moore et al.,1991 ). Neither KOH nor acetolysis treatments were used to avoid damage to dinoflagellate cysts. The resulting residue from the chemical processing was then mounted on slides with glycerin jelly.

Analysis

A variety of palynomorphs were counted; dinoflagellate cysts, Halodinium, pollen and plant spores, and foraminifer linings. Halodinium is an acritarch and freshwater tracer while foraminifera linings are an indicator of benthic productivity (Richerol et al.,

2008).

Palynomorphs were identified using an oil immersion objective with 60x magnification. Identification of pollen grains was made on the basis of published descriptions in accordance with Moore et al. (1991) and McAndrews et al. (1973).

Differentiation of Betula tree and Betula shrub was determined according to size, any grain that is less than 20 μm is counted as Betula shrub and any larger than 20 μm is

Betula tree (Leopold, 1956; Clausen 1962; Birks, 1968; Dyer, 1986). There were several depths at which the pollen count was below an desired threshold, 300 grains, and therefore have a greater uncertainty associated with them (1270 cm, 140-1522 cm, 1560-

1642 cm, 2160 cm, 2210 cm, 2340 cm, and 2360 cm).

Identification of dinoflagellate cysts were made on the basis of published descriptions in accordance with Rochon et al. (1999). Differentiation of Brigantedinium cariacoense and Brigantedinium simplex was uncertain thus they were grouped together as Brigantedinium spp. Spiniferites spp. includes all species of Spiniferites. There were several depths at which dinoflagellate cyst count was below an acceptable threshold, 100 cysts, and therefore have a greater uncertainty associated with them (850 cm, 1020 cm,

1050 cm, 1460 cm, 1560 cm, 1800 cm, and 2160 cm).

The pollen and dinoflagellate assemblages were zoned using CONISS, a constrained cluster analysis by sum-of-squares developed by Grimm (1987) and available in Psimpoll (Bennett, 2007). The taxa used for the zones are those displayed in the diagrams. Only those pollen and plant spore taxa that exceeded 5% at any one level were considered while only those dinoflagellate cyst taxa that exceeded 3% at any one level were considered.

Air temperatures were reconstructed using transfer function based upon modern analogues (Overpeck et al., 1985; Birks, 1995) using C2 software (Juggins, 2010).

Temperature reconstructions were performed for both February and August in order to represent the coldest and warmest months of the year. Pollen and climate data were taken from the North American Modern Pollen Database (Whitmore et al., 2005). The modern samples used include all lacustrine sites east of -85.0° and north of 45.0°. Pollen data from peat bogs and unknown locations were omitted, leaving a total of 651 sites used in the analysis. Bootstrap cross validation was performed in order to identify any outliers within the modern samples. The pollen sum was based on total tree, shrub and herb pollen, while the climatic transfer function was based on the taxa represented in the pollen diagram (Figure 5) excluding Sphagnum and Polypodiaceae. Sphagnum and

Polypodiaceae were excluded from the transfer function because they most likely represent long distance transport in the form of fluvial input.

Sea surface conditions (temperature, salinity, duration of seasonal sea ice cover) were reconstructed by a modern analog technique from a dinoflagellate cyst r database (de

Vernal et al., 2001; Guiot and de Vernal, 2007; R Development Core Team, 2009).

Chapter 3. Results

Palynomorph concentrations

The palynomorph concentrations varied greatly with depth. Dinoflagellate cyst concentrations ranged from 104-1,817,515 cysts cm-3 of sediment, pollen concentrations ranged from 142-1,412,384 grains per cm-3 of sediment, Halodinium ranged from 6.5-

16,725 per cm-3 of sediment and foraminifera linings ranged from 80-1,070,438 per cm-3 of sediment. Between 1,700-1,800 cm depth palynomorph concentrations peak and pollen grain concentrations average 315,719 pollen grains per cm3 of sediment, dinoflagellate cysts concentrations average 389,900 cysts per cm-3 of sediment. Between 1170-1230 cm depth palynomorph concentrations peak and pollen grain concentrations average 738,757 grains per cm-3 of sediment; dinoflagellate cysts concentrations average 847,042 cysts per cm-3 of sediment. These palynomorph peaks suggest lower sedimentation rates at those levels.

Comparison with other pollen and dinoflagellate cyst records in the area

Pollen

The Bay of Islands record has only two 14C dates, so to create an age model for core MD99-2225 comparisons of pollen and dinoflagellate cyst profile features or zone boundaries were made with surrounding records.

There are several similar trends within Bay of Islands record and surrounding records, such as at Joe’s Pond (McCarthy et al., 1995), Southwest Brook Lake (Anderson 20

and Lewis, 1992), Robinsons Pond (McCarthy et al., 1995), Northern Baie Verte

Peninsula (Dyer, 1986), and Compass Pond (Dyer, 1986) (see Table 4).Although

Northern Baie Verte Peninsula and Compass Pond are not the closest sites geographically to Bay of Islands, their pollen sequences show similar pollen zonations in the early

Holocene and they are also the only nearby records going as far back in time. Both sites had high grass-herb-Artemisia percentages within their base zones recalibrated to 14000 yr BP, for Northern Baie Verte Peninsula and 12000 yr BP, for Compass Pond which were also found within Zone 1 of the Bay of Islands core.

The assemblages within zone 2, based on low Betula percentages as well as the high percentages of grass, Cyperaceae and Salix (see Table 4) from Bay of Islands are very similar to Joe’s Pond, located on the southwest of Newfoundland, Southwest Brook

Lake, dated 11300 yr BP, Robinsons Pond, Northern Baie Verte, dated 10000-11800 yr

BP, and from Compass Pond, dated 10000-12000 yr BP. The assemblages within zone 2 of Bay of Islands are very similar to those from southwest Newfoundland lakes (Table 4) as they all show low Betula percentages, high percentages of Poaceae, Cypercaceae and

Salix. This pollen zone ranges between 10,000 and 12,000 yr BP in most records (Dyer

1986, Anderson and Lewis, 1992, McCarthy et al. 1995).

The lower section of zone 3 from Bay of Islands shows a spike in Betula around

1970 cm. This is also seen at s Joe’s Pond, Southwest Brook Lake, dated 10500 yr BP, and Robinsons Pond (McCarthy et al. 1995). The middle and upper section of zone 3 from Bay of Islands show increased percentages in tree taxa such as Betula, Pinus and

Picea, this trend is also seen at Joe’s Pond. There is a notable spike in Picea between

1870-1800 cm which is also seen at Southwest Brook Lake, and Compass Pond both dated 9500 yr BP. Anderson and Lewis (1992) and Macpherson (2005), maintain that

Picea became established in southwestern Newfoundland ~10,000 yr BP; this suggests that Picea had been growing in the vicinity of Bay of Islands and Joe’s Pond since 9,500 yr BP. Picea pollen found earlier in the record was most likely wind-blown from a more southerly site. A short-lived spike in Betula at 1700 cm is also recorded in both Joe’s

Pond and Southwest Brook Lake at 9000 yr BP. This peak is believed to represent a warming period (Macpherson 2005).

Zone 4 from Bay of Islands shows a decrease in trees and an increase in shrubs as well as high percentages of herbs, this trend is also found within the Southwest Brook

Lake record (8700 yr BP).

At the base of zone 5, at 1500 cm depth, there is a peak in Pinus pollen recorded in Bay of Islands that is also found at Joe’s Pond. This peak in Pinus most likely represents the arrival of Pinus in the area. Macpherson (2005), found Pinus at southwest

Brook Lake, about 70 km south of Bay of Islands, around 8,500 yr BP and if their analysis is correct it is perfectly reasonable for Pinus to become established in Bay of

Islands by 9,000 yr BP.

Dinoflagellate cysts

There are several trends within the Bay of Islands record that are similar to those from Cabot Strait, Esquiman channel, Anticosti channel (de Vernal et al., 1993), the

southern Labrador Sea, off the coast of Newfoundland, (de Vernal and Hillaire-Marcel,

2000) and St. Anne’s Basin (Levac et al., 2011) (Table 5).

Postglacial dinoflagellate cyst assemblages from the Gulf of St. Lawrence are fairly consistent which led de Vernal et al., (1993) to define a two part regional ecostratigraphy for the Gulf of St. Lawrence. Part 1 represents the cold late-glacial conditions of pre- 10,000 yr BP while part 2 represents the warmer later glacial period.

The assemblages that comprise part 1 are of low species diversity and mainly dominated by Brigantedinium spp. and Operculodinium centrocarpum. These low diversity assemblages are also found from 2500 to 2050cm depth in Bay of Islands. The dinoflagellate cysts assemblages from the Gulf of St. Lawrence recorded a sharp transition to a higher diversity cyst assemblage at around 10,000 yr BP, which is believed to represent the establishment of postglacial (Holocene) sea surface conditions (de Vernal et al., 1993). These changes are also recorded in Bay of Islands.

There are also several similarities between core MD99-2225 and core HU91-045-

094 (here after referred to as core 94) off the coast of Newfoundland, in the southern part of the Labrador Sea (de Vernal and Hilliare-Marcel, 2000). A spike in Brigantedinium spp. in zone 2 from Bay of Islands is also seen in core 94 dated around 12,300 yr BP.

High concentrations of Operculodinium centrocarpum, in Bay of Islands’ dino zone 3, are also seen in core 94, dated around 11,140 yr BP. High concentrations of

Brigantedinium spp., and a peak in Impagidinium aculeatum that characterize Bay of

Islands dinoflagellate cyst zone 4 are also seen in core 94 around 11,140-10,020 yr BP.

Consistent increases in Bitectatodinium tepikiense and Islandinium minutum as well as decrease in concentrations of Brigantedinium spp. within dinoflagellate cyst zone

5 are seen within the upper section of core 94. Bay of Islands dino zone 5 shares a few similar trends with the dinoflagellate cyst assemblages from core 84-011-12 from St.

Anne’s Basin, hereafter referred to as core 84 (Levac et al., 2011). Two spikes in

Pentapharsodinium dalei are also recorded in core 84 and dated at approximately 8,800 and 8,100 yr BP, respectively. Above this, spikes in Brigantedinium spp. and O. centrocarpum are seen around 7,800 and 7,700 yr BP, respectively, within core 84. In the middle of zone 5, increases in Selenopemphix quanta and Nematosphaeropsis labyrinthus also occur in core 84 between roughly 8,800 and 7,100 yr BP. Near the top of zone 5,

Bitectatodinium tepikiense, Brigantedinium spp. and Operculodinium centrocarpum spike; this also occurs in core 84 at approximately 7,100 yr BP 6,900 yr BP and a spike in

Operculodinium centrocarpum is also seen in core 84 at approximately 6,700 yr BP.

Pollen Assemblage Zones

The pollen assemblages from the Bay of Islands core is divided into five zones

(Figure 4).

Pollen assemblage Zone 1: (2450-2500 cm). Pinus is present in high percentages, but shrub pollen, particularly from Betula, Alnus and Salix is important in this zone. Pollen from sedges also is important. Other taxa include Ambrosia and Sphagnum.

Pollen assemblage Zone 2: (2090-2450 cm). As in zone 1, shrub and sedge pollen are abundant, but percentages of Pinus pollen decrease towards the top of the zone. Shrubs including Betula, Alnus, and Salix. Ambrosia, Sphagnum, and Polypodiaceae are present in small percentages, around 15% each.

Pollen assemblage Zone 3: (1650-2090 cm). Shrub and sedge pollen continues to be important and the first appearance of Fraxinus occurs. Sphagnum and grass are also present in moderate proportions. Betula replaces Alnus as the dominant shrub species from zone 2. Towards the top of zone 3 there is an increase in tree pollen from Betula,

Picea, Abies, and Acer. Between 1700 and 1800 cm both tree Betula and Picea pollen percentages peak, concurrent with a peak in the total palynomorph concentrations at this depth (Figure 3).

Pollen assemblage Zone 4: (1550-1650 cm). This zone is essentially barren with a maximum of 71 pollen and spore grains. No Abies, Acer, Tsuga or tree Betula were counted. Pinus, Salix, Ericaceae, Alnus, Cyperaceae and Poaceae do appear in this zone.

Pollen assemblage Zone 5: (1000-1550 cm). The abundance of pollen from grass, sedge,

Salix and Ericaceae overall, is lower in zone 5 compared to previous zones. There is an increase in the more temperate trees and shrub species such as Acer and Fraxinus.

Towards the upper section of zone 5 percentages of Betula pollen display a reciprocal pattern from pollen of conifers. Between 1,040 and 1,100 cm depth, Betula peaks, this spike is also found within the total pollen concentrations at this depth.

Dinoflagellate cyst Assemblage Zones

The dinoflagellate cyst diagram from Bay of Islands core is clustered into five assemblage zones (Figure 5).

Dinoflagellate cyst assemblage Zone 1 (2470-2500 cm) Zone 1 assemblages are represented by Pentapharsodinium dalei and Brigantedinium spp. Scrippsiella trifida, and Operculodinium centrocarpum are present in small but noticeable proportions, 15 and 5% respectively.

Dinoflagellate cyst assemblage Zone 2 (2030-2470 cm) The transition from zone 1 to zone 2 is marked by a shift in from dominance P. dalei to Brigantedinium spp.

Percentages of P. dalei decrease towards the top of the zone. Other taxa present are

Bitectatodinium tepikiense, S. trifida, Islandinium minutum, O. centrocarpum,

Selenopemphix quanta, and Spiniferites spp. Towards the top section of the zone S. trifida, and I. minutum are represented in high proportions and Ataxiodinium choanum make their first appearance in the assemblages.

Dinoflagellate cyst assemblage Zone 3 (1630-2030 cm) The assemblages in the lower section of zone 3 are represented by S. trifida, as well as O. centrocarpum and B. tepikiense which both experience a significant increase compared with zones 1 and 2,

While proportions of Brigantedinium spp., I. minutum, and P. dalei, are much reduced.

Ataxiodinium choanum is consistently present. The middle section of zone 3 shows some notable changes such as a significant decrease in S. trifida and an increase in Spiniferites spp.

Dinoflagellate cyst assemblage Zone 4 (1450-1630 cm) The assemblages in the lower section of zone 4 are represented by Brigantedinium spp. and O. centrocarpum both of which shows large variations in percentages. Other taxa present are P. dalei, I. minutum, and Spiniferites spp. Scrippsiella trifida and P. dalei decrease or disappear in the lower part of the zone but increase near the top, while B. tepikiense and O. centrocarpum show the opposite trend. Nematosphaeropsis labyrinthus makes its first appearance within the lower section of zone 4.

Dinoflagellate cyst assemblage Zone 5 (1000-1450) The assemblages in the lower section of zone 5 are represented by P. dalei, O. centrocarpum and Spiniferites spp. while

Brigantedinium spp. experiences a notable decrease from zone 4. Bitectatodinium tepikiense, S. trifida, Spiniferites spp., I. minutum, A. choanum, N. labyrinthus, and S. quanta are all present within the lower section of zone 5 in notable proportions.

Operculodinium centrocarpum, and S. trifida peak ~ 1000 cm depth. At the very top (850 cm) Brigantedinium spp. O. centrocarpum and I. minutum dominate the assemblage.

Reconstruction of sea surface conditions

Figure 6 shows the reconstructions of sea surface conditions. Overall, reconstructed salinity varies little from present day which averages 32 psu in February and 30 psu in August. August sea surface temperatures were lower than present,

16.96°C, for most of the record.

From 14,000 to 12,000 yr BP (2,400-2,100 cm depth, dinoflagellate cyst zone 2 and 3), sea surface temperatures were much lower than today. At the base of the record,

August SST is around 12°C and between 2400-2100 cm August SST oscillates between

0-5°C. Between 2400-2500 cm, February SST is about the same as today, but drops to -

2°C between 2400-2100 cm. The average ice cover during this period was 7.4 months yr-

1which is substantially longer than the 2.5 -3 months yr-1duration that occurs presently in

Bay of Islands.

Around 2100 cm (ca. 12,200 yr BP) SST in both February and August increased markedly and the sea ice cover duration reduced. Between 2100-1800 cm temperatures were about 2°C higher than the modern temperature in February and about 3°C higher than the modern temperature in August. Sea surface salinities were close to modern conditions. The average duration of ice cover during this period decreased to roughly 2.7 months yr-1 slightly longer than today. Between 1800-1400 cm SST fluctuated greatly and remained generally under the present day averages.

Between ca 10,000 and 7,000 (1450-1050 cm depth, dinoflagellate cyst zone 5), both temperatures and salinity showed considerable variability. August temperatures dropped by 1°C during two intervals: from 1550-1450 cm and from 1400-1300 cm. These same intervals are characterized by increased duration of ice cover.

Reconstructed air temperatures

The reconstructed air temperatures for February over the studied interval were consistently lower than today's average of -7°C. February temperatures ranged from -

23.06 °C to -9.5°C with an average temperature of -16°C (Figure 7.). The reconstructed air temperatures for August overlapped with today's August average of 16.9°C. Past

August temperatures ranged from 5.21°C to 17.24 °C with an average temperature of

13.4°C.

Reconstructed air temperature at the base of the Bay of Islands core (ca. 14,000 yr

BP) represent postglacial conditions, with February temperatures much colder than today, oscillating between -20 and -15 °C, while August temperatures are closer to the present- day average, oscillating between 10 and 15°C. Following this, and still within the interval representing pollen zone 1 and 2, there is a sharp drop in temperatures. Between 2100 and

2300 cm (ca. 11,500-12,500 yr BP) February temperatures remain between -23 and -19

°C and August temperatures remain between 6 and 12 °C.

There is a marked increase in temperature around 12,000 yr BP (2000 cm).

Between 12,000 and 10,000 yr BP (pollen zone 3 and lower part of zone 2).. With values usually between 15 and 16 °C, August temperatures are close to the present-day average.

Both February and August temperatures show large oscillations, with noteworthy drops in temperature around 2000 cm, 1900 cm, and 1800 cm.

The interval between 1650 and 1450 cm (ca. 9,250-9,750 yr BP) (interval corresponding to pollen zone 4 and lower pollen zone 5), is characterized by large drops

in temperatures with February temperatures falling as low as -22 °C. August temperatures drop below 10 °C in most of the interval and to 5°C at 1600 cm.

In the interval between 1450-1000 cm (ca.10,000 and 7,000 yr BP; pollen zone 5) the average February temperature was -15.2°C, still colder than present day averages but with an overall increasing trends. Average August temperatures increased slightly to

14.7°C, close to current conditions. While August temperatures show very little variation during this interval, February temperatures are variable throughout this period, with a sharper drop below -17°C between 1160 and 1280 cm.

Chapter 4. Discussion

Younger Dryas

The Younger Dryas is evident within both the pollen and dinoflagellate record in

Bay of Islands. There were high percentages of the shrubs Alnus and Salix, sedge, and grass pollen. The Pinus between 2,350-2,000 cm, within zone 2 and bottom of zone 3 are assumed to be from long distance transport (Figure 5) as Macpherson (1995) has evidence that Pinus did not arrive in southwestern Newfoundland until ~8,500 yr BP.

Similar assemblages are reported at Robinsons Pond, Southwest Brook Lake, Joe’s Pond

(McCarthy et al., 1995; Anderson and Lewis, 1992) and at Leading Tickles (Macpherson and Anderson, 1985) (Table 4). Reconstructions based on pollen show that air temperature around Bay of Islands dropped by 7°C at the start of the Younger Dryas and remained between about 6-8°C below the air temperature before and after the event.

February air temperatures dropped by 7°C at the onset of the event which is consistent with the 6-7°C decreased noted by Mayle and Cwynar (1995). In the dinoflagellate cyst record from the Bay of Islands, the Younger Dryas is characterized by very low species diversity, and unique assemblages dominated by mostly S. trifida and Brigantedinium spp. This assemblage is present between 2,350 and 2,000 cm (within dinoflagellate cyst zone 2 and the bottom of zone 3). August sea surface temperature and salinity dropped around 10°C and to 28 psu, from 30 psu, at the onset of the Younger Dryas, and the seasonal duration of sea ice increased from 2-3 months to 10 months yr-1.Air temperature seems to lag SST by 100 years at the beginning of the Younger Dryas. This can be explained by the rapid turnover rate of dinoflagellate cysts which are quick to respond to

environmental change. The climate possibly changed too rapidly for vegetation to respond; slow growing tundra plants require more time to react than more rapidly reproducing vegetation in warmer climate. This lag in vegetation response also could be due to the lingering presence of ice, land was not available for vegetation to return.

At the end of the Younger Dryas a lag between the dinoflagellate cysts and vegetation’s response also occurred. This lag could be due to the extended presence of ice even after the temperatures had changed that inhibited the vegetation’s response.

Although there was a lag time experienced between the response of the dinoflagellate cysts and vegetation, they were still able to respond to the changing conditions.

The impact of the Younger Dryas has also been reported outside of

Newfoundland at several sites in eastern North America. Levesque et al., (1997) report simultaneous responses to the Younger Dryas in the pollen record from five 5 lakes located between central Nova Scotia and Maine. The magnitude of the cooling, based on palaeoclimate reconstructions from fossil assemblages of aquatic midge larvae from the 5 lakes, in Eastern North America was between 6 and 20°C, with the absolute cooling increasing southward (Levesque et al., 1997). The cooling throughout eastern North

America was so significant that is caused vegetation succession reversal. Mayle and

Cwynar (1995) analyzed cores from lakes in the Canadian Maritimes and found that the boreal forest was replaced by shrub tundra in southern New Brunswick and central mainland Nova Scotia and the shrub tundra was replaced by herb tundra in central New

Brunswick, northern Nova Scotia and Newfoundland. Mott et al. (1986) had previously noted vegetation changes in 14 lakes throughout Maritime Canada.

The Picea and Populus forest present in southern New Brunswick and southern

Nova Scotia and the shrub-herb vegetation present in northern New Brunswick and northern Nova Scotia before the Younger Dryas returned to a more open shrub-herb vegetation around 11,000 yr BP. Williams et al. (2002) reconstructed vegetation changes at Splan Pond, New Brunswick (Levesque et al., 1997), and noticed decreases in Picea and other arboreal pollen and increases in Cyperaceae and Poaceae during the Younger

Dryas. The termination of the Younger Dryas event and start of the PBO was also recorded as a significant transition period within pollen records.

Regionally, few records of sea surface conditions extend as far back as the

Younger Dryas (Table 4). In Cabot Strait, de Vernal et al., (1993) reports assemblages dominated by Brigantedinium spp. before ca.10,000 year BP. These would be associated with relatively cold surface waters and extensive sea ice cover (approximately 8 months yr-1). Pollen assemblages from sediments of Cabot Strait reported by Giroux (1990) for the same time period were typical of open tundra vegetation. Sea surface reconstructions for the period between ca. 11,800-10,000 yrs BP in Cabot Strait show a period of low sea salinity, which they believe corresponded with the Younger Dryas event (de Vernal et al.,

1993). Within the southern part of the Labrador Sea de Vernal and Hilliare-Marcel

(2000) also found evidence of the Younger Dryas. They found dinoflagellate cyst assemblages with low species diversity and high percentages of Brigantedinium spp.

Reconstructions of sea surface conditions show low salinity and extensive sea ice cover of up to 12 months yr-1 as well as cold sea surface temperatures, however, the very low resolution of this core prevents a precise dating for the event, and high bioturbation values in this area (Anderson, 2001) would certainly smooth the signal.The low species 33

diversity, low salinity and extended period of ice cover reflect the large volumes of melt water that were released into the ocean from the waning Laurentide ice sheet at the start of the Younger Dryas. At other sites in Cabot Strait de Vernal et al. (1996) interpreted dinoflagellate cyst assemblages to indicate the influx of meltwater at the time of the

Younger Dryas along the path proposed by Broecker et al. (1989) and Carlson (2010).

Carlson (2010) shows evidence of meltwater drainage through the St. Lawrence River..

They found evidence of reduced melt-water runoff indicated by an increase in salinity, cooler temperatures and longer duration of seasonal sea-ice cover during the Younger

Dryas, between ca. 10,800-10,300 yr BP. However, de Vernal et al. (1996) did note that they found melt-water pulses both before and after the Younger Dryas, as early as ca.11,700 yr BP and as late as ca. 10,100 yr BP. They believe the melt-water pulse recorded prior to the Younger Dryas did not affect the sea surface salinity, therefore, did not support a weakened North Atlantic thermohaline circulation leading to cooling

(Broecker et al., 1989).

The termination of the Younger Dryas event and start of the PBO was also recorded as a significant transition period within marine records. In the Cabot Strait record de Vernal et al. (1993) noted a sharp transition at ~10,000 yr BP, indicated by an increase in Gonyaulacales, Bitectatodinium tepikiense and Alexandrium tamarense, which are associated with warmer temperatures, higher salinity and shorter seasonal ice- cover (approximately 2 months yr-1) - conditions similar to those of today. This transition marks the termination of the Younger Dryas event. Bay of Islands also records an increase in B. tepikiense at the termination of the cold event as well as increases in other warmer temperature taxa. Changes in δ18O of deep sea coral from Orphan Knoll in the 34

northwest Atlantic Ocean were also seen with the initiation of the Younger Dryas, suggesting profound changes in intermediate water circulation at this period (Smith et al.,

1997).

Preboreal Oscillation (PBO)

The PBO’s impact was regionally variable around Eastern North America; some records reflect the event while others did not (Anderson et al., 2007). Because the event was so short lived, most likely those sites bordering ecotones were sensitive to the change. Today, Bay of Islands is located on the boundary between the tundra and boreal forest ecotones and as a result the event was recorded within the pollen record. A cooling trend is observed within the pollen record from Bay of Islands, during which air temperatures averaged -22°C in February and 5°C in August, lower that modern average temperatures of -7.2°C in February and 16.9°C in August (Environment Canada 2011).

Temperatures dropped 9°C in August and 3°C in February at the start of the PBO.

However, because of low pollen counts within pollen zone 4 higher counts are needed to be conclusive. Several scenarios could explain the low pollen counts within this interval.

The most plausible scenario is the lower pollen production that is characteristic the type of vegetation growing in the region during the PBO. Also, there is the possibility that the

Humber River and the Bay were covered in ice at that time which inhibited pollen input to the sediments. The sediments in this interval are laminated silty clay which could be an indicator for longer sea ice cover in the Bay. The ice would eliminate sea surface currents allowing the finer silt and clay sediments to become deposited.

The PBO event was recorded within the dinoflagellate cyst record. During the

PBO sea surface temperatures dropped as much as 15°C in August and 1°C in February, sea surface salinity decreased during this period and seasonal sea ice cover increased up to 9 months yr-1. The change noted in the dinoflagellate cyst record was also significant; there was an increase in Brigantedinium spp., and S.quanta while most other species decreased during this period. The low species diversity and assemblage domination of

Brigantedinium spp. suggests colder temperatures.

Because there was insufficient pollen data collected, conclusions cannot be made about the possible lag time between the dinoflagellate cyst and vegetation response, both at the beginning and end of the PBO.

The PBO was also recorded by Anderson et al., (2007) at several sites around the

Gulf of St. Lawrence. In western and southwestern Newfoundland Picea was replaced by shrub tundra at the beginning of the cold period. Anderson et al., (2007) attributed the decrease in Picea and increase in Betula shrubs and trees to cooling around 11,026 ±311

14C cal yr BP. The PBO was also recorded at Upper South Brand Pond, in Maine where

Anderson et al., (1986) noted increased percentages in Picea pollen and decreases in

Pinus interpreted as a response to cooling during the PBO. The PBO is recorded within the Greenland Ice core records (Van Der Plicht et al., 2004) as a period of decreased snow accumulation resulting from colder temperatures and decreased available water vapour. Van Der Plicht et al., (2004) believe the cooling is a result of the large fresh meltwater pulse from the melting ice sheets as also proposed by Fisher et al., (2002) and

Broecker et al., (1989).

8.2 ka event

The response to the 8.2 ka event was variable throughout eastern North America; some sites recorded it while others did not. The magnitude of cooling in Eastern North

America ranged between 1.5 and 3°C at both marine and terrestrial sites (Barber et al.,

1999). In Bay of Islands the 8.2 ka is recorded as a decrease in tree pollen and an increase in shrub pollen during this period as well as an increase in cold temperature dinoflagellate species. Sea surface temperatures decreased by about 1°C in February and 7°C in August.

Although sea surface temperatures decreased during the 8.2 ka event, there was no change in the sea surface salinity in the Bay during this period. Although the meltwater released by Lake Agassiz was responsible for an significant important cooling significant enough to cause changes in the vegetation, the melt water path along the slope of the

Labrador shelf (Levac et al. 2011) could explain that no decrease in salinity is recorded in

Bay of Islands at that time. The air temperature seems to lag SST by 50 years at the beginning of the 8.2 ka event. The immediate decrease in SST is to be expected due to the dinoflagellate cysts rapid response to changing environments. Because the 8.2 ka event occurred so rapidly the more thermophilous vegetation died and tundra vegetation took longer to respond. At the end of the 8.2 ka event the lag in vegetation response was slightly longer than at the beginning. This extended lag time could be due to the prolonged frozen soil inhibiting vegetation succession.

The 8.2 ka event was more clearly seen within marine records and the Greenland ice core records. Thomas et al. (2006) observed a 160.5 year cold period within four high

resolution records from the central Greenland ice cap, during which the δ18O values indicate decreases in snow fall accumulation and a drop in temperature between 4-8°C during the cold event. Keigwin and Jones (1995) interpreted an increase of the cold- adapted foraminifer Neogloboquadrina pachyderma with low δ18O on the continental margin of Nova Scotia to represent cooler conditions with freshening due to an input of fresh meltwater from Lake Agassiz. On the Labrador Shelf Andrews et al. (1999) found a peak in a ‘cold fresh’ benthic indicator foraminifer. Bond et al. (1997) found a high percentage of N. pachyderma, which they suggest represents an abrupt and brief onset of cooling from two cores off the south west coast of Norway in the North Atlantic. Within cores from St. Anne’s Basin and Notre Dame Channel, Levac et al. (2011) reconstructed decreased sea surface salinities and a slight decrease in sea surface temperatures (2-3°C) from detrital carbonate layers corresponding to the 8.2 ka event. These reconstructions show that the final Lake Agassiz drainage, dated 8.33 ka with an error range of 8.15-8.48 ka (Lewis et al., 2009), significantly influenced sea surface conditions along the eastern

Canadian shelf. De Vernal et al. (1993) also found evidence of a cooler period between ca

9000-8000 yr BP from the dinoflagellate cyst record from Cabot Strait and Anticosti

Channel, but they could not verify that the cool period was caused by the meltwater drainage from Lake Agassiz. One reason for this is that the location of the cores might not have been on the meltwater drainage path (Levac et al., 2011).

Locally, Anderson and Lewis (1992) noted decreased percentages in Picea at the time of the 8.2 yr BP event and an increase in shrub Betula, which maintained dominance for 200-400 years (Anderson and Lewis, 1992), these vegetation changes were also recorded within the Bay of Islands record. Daley et al. (2009) found evidence for an 38

abrupt cooling event at ca. 8350 yr BP that lasted between 150-200 years from

δ18Ocellulose from Sphagnum mosses retrieved from a cored peatland on the east coast of Newfoundland.

Chapter 5. Summary and Conclusions

The pollen and dinoflagellate cyst assemblages from the Bay of Islands recorded shifts in vegetation and dinoflagellate cysts as a response to the Younger Dryas, PBO and

8.2 ka climatic events. This study is the first to provide reconstructions of both air temperature and SST for these climatic events that can be directly correlated, since they both come from the same core (and the same samples). For the Younger Dryas and 8.2 ka event, changes in vegetation lagged behind the changes observed in dinoflagellate cysts assemblages, both at the start of the cold events as well as during the recovery. Which means dinoflagellate cyst assemblages are more sensitive to changes in temperatures and are a better tool to detect climatic events.

This study also is the first to provide reconstructions of SST for the Younger

Dryas for the western North Atlantic Ocean. The Younger Dryas was recorded as a 10°C drop in sea surface temperature, a decrease in sea surface salinity as well as an increase in seasonal ice cover from 2-3 months yr-1to 10 months yr-1. The dinoflagellate cysts responded to the Younger Dryas with very low species diversity, dominated mainly by

Brigantedinium spp. The vegetation responded to the cooling with increased abundance of shrubs and grasses, which is indicative of a tundra vegetation that lagged by roughly

100 years relative to the shift in dinoflagellate cyst assemblages. The PBO was recorded as a 11°C drop in sea surface temperature, a slight decrease in sea surface salinity as well as an increase in seasonal ice cover from 2-3 months yr-1to 9 months yr-1. The dinoflagellate cysts responded with very low species diversity, mainly dominated by

Brigantedinium spp. The very low pollen counts in this interval prevents assessment of vegetation response but suggests extensive glacial ice

The 8.2ka event was recorded as a 10°C drop in sea surface temperature and an increase in seasonal ice cover from 2-3 months yr-1 to 9 months yr-1.The dinoflagellate cysts responded with an increase in cold water species such as Brigantedinium spp., I. minutum and P. dalei. The vegetation responded, roughly 50 years later, with an increase in shrubs and a decrease in trees, which represents a shift to a tundra vegetation.

It has been questioned what the vegetation response will be to the abrupt climate changes that are expected in the future. Based on what is seen in the pollen record for the

Younger Dryas and 8.2 ka event, although there is a lag associated with the vegetation response, the thermophilous vegetation present at the beginning of the event dies off and the vegetation more suited to the colder climate grows, and the response might be faster in warmer regions.To draw more accurate conclusions additional high resolution marine records should be used so that the marine signal can be used to validate the terrestrial signal. More high resolution marine cores, with sufficient 14C dates, also would be useful because they allow for a clear connection between what was occurring on land as well as in the sea at the same time which would allow for a better understanding of the adaptability, response and lag times associated with different vegetation types. In order for this to be reliable however, marine environment modern analogues for pollen need to be developed.

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Table 1. Dominant vegetation presently found within the western Newfoundland ecoregion (Damman, 1983; Davis, 2001).

Family Species Aceraceae Acer rubrum Aceraceae Acer spicatum Betulaceae Alnus rugosa Betulaceae Betula alleghaniensis Betulaceae Betula lutea Betulaceae Betula papyrifera Cornaceae Cornus solonifera Hylocomiaceae Hylocomium splendens Oleaceae Fraxinus nigra Pinaceae Abies balsamea Pinaceae Picea glauca Pinaceae Picea mariana Pinaceae Pinus strobus Salicaceae Populus tremuloides Salicaceae Salix spp.

Table 2. Radiocarbon dates for core MD99-2225, Bay of Islands. Calibrated Calibrated Type Age 1σ Age 2σ of C14 range cal range cal Depth Lab no. Sample age ∆R BP BP 1440 TO- Hinged 9120± 150± 9520 to 9970 to cm 8460 bivalve 90 42 9790 10000

2166 TO- Hinged 10980± 150± 12460 to 12000 to cm 8461 bivalve 80 42 12530 12580

Table 3. Summary of radiocarbon dates from sites near Bay of Islands. Cores from St. Anne’s Basin (Levac, 2002), Bay of Islands (levac, 2002), Leading Tickles (Macpherson and Anderson, 1985), Compass Pond (Dyer, 1986), Southwest Brook Lake (Anderson and Lewis, 1992), Joes Pond (McCarthy et al., 1995) and Robinsons Pond (McCarthy et al., 1995). Age Age 1σ 2σ This Conventional range range Corresponding study Lab Depth 14C Age, Dated (cal yr (cal yr Bay of Islands depth Location no. (cm) years B.P. material ∆R BP)` BP) zone (cm) 2729 2687 St. Anne's TO- to to Basin 6232 65-67 3100 ± 60 Foraminifera 83±14 2850 2945 290 2851 2765 Bay of TO- to to Islands 8459 341 3310±60 Tree bark 150±42 3057 3161 341

Leading GSC- 210- 4778- 5012- Tickles 4107 215 4200±110 Bulk 4855 5036 600 5278 4960 Compass GSC- 295- to to Pond 3903 300 4690±160 Bulk 5602 5733 660 7010 6903 Compass GSC- 395- to to Pond 3902 400 6280±120 Bulk 7321 7428 5 950

Leading GSC- 265- 6656- 7095- Tickles 4085 270 5960±120 Bulk 6953 7157 5 8229 8146 St. Anne's TO- 640- to to Basin 6233 650 7930±80 Foraminifera 83±14 8386 8490 5 1200 Table 3 continued

SW 9291 9076 Brook GSC- 314- to to Lake 5041 320 8550±220 Bulk 9894 10176 5 1300 9460 9550 Compass 435- to to Pond 435 440 8310±140 Bulk 9140 8820 5 1350 9518 9974 Bay of TO- Hinged to to Islands 8460 1440 9120±90 bivalve 150±42 9787 10002 5 1440 11663 11094 Compass GSC- 495- to to Pond 3898 500 9950±150 Bulk 11708 12051 3 1750 10218 11870 GX- 422- to to Joes Pond 9963 425 9445±380 Bulk 11246 11965 3 1800 12265 10663 GX- 490- to to Joes Pond 9964 495 10130±375 Bulk 12381 12688 2 1900

Leading GSC- 380- 10649- 11514- Tickles 4183 385 9600±230 Bulk 11224 11613 3 1930 12521 11594 Robinsons GX- 408- to to Pond 9965 417 11300±620 Bulk 13954 14987 2 1950 SW 12834 12689 Brook GSC- 383- to to Lake 4631 386 11100±120 Bulk 13127 13230 2 1950

Table 3 continued

Leading GSC- 404- 12372- 11973- Tickles 3610 415 10500±140 Bulk 12577 12675 3 1970 12458 12000 Bay of TO- Hinged to to Islands 8461 2166 10980±80 bivalve 150±42 12526 12579 2 2166 13385 13221 Compass GSC- 540- to to Pond 3891 545 11700±180 Bulk 13739 13920 4 2300

Leading GSC- 433- 15508- 14993- Tickles 3608 443 13200±300 Bulk 16642 16890 1 2480

Table 4. Pollen assemblage comparisons from sites near Bay of Islands. Cores from Bay of Islands (this study), Joes Pond (McCarthy et al., 1995), Southwest Brook Lake (Anderson and Lewis, 1992), Robinsons Pond (McCarthy et al., 1995), Northern Baie Verte Peninsula (Dyer, 1986) and Compass Pond (Dyer, 1986).

Southwest Robinsons Northern Baie Compass Bay of Islands Joes Pond Brooke Lake Pond Verte Peninsula Pond depth age (ka- (cm) taxa characteristics BP) 1500 spike in Pinus 8500 1700 Spike in Betula 9000 9000 1820 Spike in Picea 9500 9500 8500 -10000 9500 1970 Spike in tree Betula 10,500 10500 10,500

Decreases in Betula, high % of 2090- grass, cyperaceae Comment [GC1]: see previous 2450 and salix 11,300 11300 11,500 10000-11800 10000-12000 comments about capitalization, italics, mixing scientific and common names of 2450- Shrub-Grass- families 2500 Artemesia 11800 12000

Table 5. Fluctuations of dinoflagellate cyst taxa common to Bay of Islands (this study) and nearby sites at St. Anne’s Basin (Levac, 2002), and the Southern Labrador Sea (de Vernal et al., 1993). St. Anne's Southern Labrador Bay of Islands Basin Sea depth (cm) taxa characteristic age (BP) 1000 Spike in O. centrocarpum 6,700

1030 Spike in Brigantedinium spp. 6,900

1040 Spike in B. tepikiense 7,100

1050 Spike in S. ramosus 7,400

1090 Spike in P. dalei 7,500

1180 Spike in O. centrocarpum 7,700

1190 Spike in Brigantedinium spp. 7,800

1210 Spike in P. dalei 8,100

1310 Spike in P. dalei 8,800 170- 180 cm 1380 Spike in O. centrocarpum (10,020-10,030) 175 cm 1500 Spike in I. aculeatum (10,000) 1480- S 180-190 1500 decrease in O. centrocarpum (11,120- 10,020) 1401- 180-190 cm 1610 Spike in Brigantedinium spp. (11,120-10,020 ) 190-195 cm 1581 Spike in O. centrocarpum (12,000-11,140 ) 200 cm 2180 Spike in Brigantedinium spp. (12,300 )