Lateglacial Climate Change and Chronology at Lurga, Western Ireland - Derived from Multiproxy and Microtephra Analysis

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Lateglacial climate change and chronology at Lurga, western Ireland - derived from multiproxy and microtephra analysis -

Master thesis

M.C.H. Duijkers, 2009 Studentnr: 0314056 Supervised by Drs N. van Asch, Dr. W.Z. Hoek Department of Physical Geography Faculty of Geosciences Utrecht University Contents

Pagea Figures 3

1. Introduction 5

2. Lateglacial Climate 7 2.1 Ice cores 7 2.2 Marine cores 11 2.3 The Lateglacial in northwest-Europe 14 2.4 The Lateglacial in Ireland 23

3. Study Area 31

4. Methods 33 4.1 Climate reconstruction 33 4.1.1 Lithostratigraphy 33 4.1.2 Loss on ignition 33 4.1.3 Isotope analysis 35

4.2 Chronological Framework 40 4.2.1 Tephra 40

5. Results 47 5.1 Climate reconstruction 47 5.1.1 Lithostratigraphy 47 5.1.2 Loss on ignition 50 5.1.3 Isotope analysis 52

5.2 Chronological Framework 56 5.2.1 Tephra 56

6. Discussion 58 6.1 Organic matter content 58 6.2 Isotope analysis 59 6.3 Tephra 55

6.4 Climate reconstruction 66

7. Conclusions 69

Acknowledgements 70

References 71

Figures

Fig. 1: δ18O curve for the Lateglacial derived from the NGRIP ice core 5 Fig. 2: The last 150.000 years in the Quaternary period. The Lateglacial forms the transition from the Late 7 Weichselian to the Holocene. Fig. 3: Map of Greenland showing sites where ice cores were retrieved. 8 Fig. 4: δ18O curve for the Lateglacial derived from the NGRIP ice core). GICC05 chronology events are 9 indicated. Fig.5: Location of marine cores MD95-2006 and NA 87-22 11 Fig. 6: Correlation of grain parameters and planktonic foraminifera from marine core at Barra Fan to NGRIP 12 ice core oxygen isotope record. Fig. 7: Sea surface temperature and salinity estimates for core NA 87-22 off Ireland. Results are compared to 13 NGRIP δ18O record. Fig. 8: Nomenclature differences for the Lateglacial period on the British Isles and in Continental North-west 14 Europe. Fig. 9: Temperature curves for different regions in northwest Europe. The curves for Britain and the 16 Netherlands are based on fossil coleopteran records. The curves for Ireland and northwest Germany are based on palaeobotanical data. Fig. 10: Chironomid-inferred mean July air temperature record for Whitrig Bog, southeast Scotland compared 18 to NGRIP ice core oxygen isotope record. Fig 11: Chironomid-inferred mean July air temperature record for Hawes Water, northwest England (A) 19 compared to oxygen isotope record from the same core (B) and NGRIP oxygen isotope record (C). Fig. 12: Vegetation development for different regions in northwest Europe during the Lateglacial. 21 Fig. 13: Development of geomorphology and soils during the Lateglacial in northwest Europe. 22 Fig. 14: Glaciation limits during recent glaciations in Britain. Sites referred to in the text are indicated in the 24 figure. Fig. 15: LOI and Oxygen isotope record for Red Bog compared to the oxygen isotope record from NGRIP ice 25 core. Fig. 16: Oxygen isotope record for Lough Gur compared to the oxygen isotope record from NGRIP ice core. 26 Fig.17: Percentage pollen curves for Betula and Juniperus, and δ18O, erosion, and July temperature values 27 compared to NGRIP oxygen isotope record. 18 Fig. 18: Lithology (A), weight percent total calcite (B), weight percent total organic matter (C), δ Ocalcite values 28 vs. age in cal yr BP with significant climate events labeled in brackets and a 50-year moving average in grey 18 inllustrating first order trends, and the NGRIP δ Oice record for comparison. Fig. 19. Map or Ireland with the study site indicated (A), topographical map of study site Lurga (B) and a 31 detailed map of study site Lurga (C) Fig. 20: Bo-plot of SiO2 and K2O concentrations in tephras derived from the main European ash provinces 41 Fig. 21: Basaltic tephra shards of the Vedde ash from Loch Ashik, Scotland, exhibiting examples of chemical 42 alternations: (A) the formation of a leached Si gel layer, (B) pitting erosion through cation removal, (C) preferential leakage of particular phases of the glass and recrystalisation of leached products onto the glass surface. Fig. 22: Sample loss due to NaOH preparation on 20 paired samples of glass extracted from a bulk sample of 43 the Baia Averno tephra. Fig. 23: δ18O curve for the Lateglacial derived from the NGRIP ice core with expected microtephra layers in 46 Ireland Fig. 24: Coring transect Lurga, western Ireland. 49 Fig. 25: a) lithology of the studied sediment core. b) the organic matter content in percentages derived from 51 the loi-method. The red lines define the different sample series which were placed in the oven for combustion. c) the organic matter content corrected for carbonates partially burned in the oven dependent on their position. After correction many values were below zero, as this is not possible these were set to zero percent. Fig. 26: Percentage of calcite burned after combustion at 550 ºC for four hours. The top of the figure is the 51 back of the furnace and the bottom part of the figure corresponds with the front of the furnace. Fig. 27: Measured calciumcarbonate percentages (LU-A) 52 Fig. 28: Oxygen isotope result for the study site Lurga 53 Fig. 29: Wigglematched δ18O curve of the study site (from a depth of 770 cm) to the NGRIP δ18O curve 54 Fig. 30: Estimated accumulation rate at the study site compared to the δ18O curve of the study site 54 Fig. 31: LU-A δ13C results compared to LU-A δ18O results 55 Fig. 32: Graph with potential tephra shards found in core LU-A 56 Fig. 33: Shard population found at the study site in core LU-A at a depth of 750 cm 57 Fig. 34: Glass shard found during optical identification of tephra shards, thought to originate from (with glass 57 shards) contaminated cover glass. Fig. 35: Organic matter content from the study site (LU-A) compared to LOI results from a previous research 59 at Lurga Fig. 36: Carbonate percentages (of this study) compared to the NGRIP δ18O and to a carbonate curve from a 60 study at Tory Hill Fig. 37: Oxygen isotope result compared to the NGRIP δ18O curve 61 Fig. 38: Oxygen isotope results for a) Red Bog b) Lough Gur (Ahlberg et al., 1996) c) Lurga (this study) and d) 62 the NGRIP ice core Fig. 39: Sea surface and salinity estimates for core NA 87-22 off the west-coast of Ireland. Results are 62 3 compared to the δ18O result of this study. Fig. 40: Composite pollen curves from a previous research at the study site Lurga 64 Fig. 41: δ13C values of the study site compared to a δ13C record from a previous study at Tory Hill 65 Fig. 42: Multiproxy analysis on sediment core LU-A at the study site Lurga. Results include lithology, 67 carbonates, oxygen and carbon isotope ratios Fig. 43: LU-A δ18O results compared to results for a chronomid detrended correspondence analysis (DCA in 68 standard deviation units) on the LU-A sediment core

1. Introduction

The Quaternary is a geological period which is characterized by the alternation of glacial and interglacial periods. The transition from the last glacial to the present interglacial (the Lateglacial) was a period of relative rapid climate fluctuations. The most important drivers of the Lateglacial climate are variations in the size of ice sheets, seasonal insolation, and greenhouse gas concentrations (Ruddiman, 2001). The Lateglacial climate fluctuations are recorded in ice cores from the Greenland Ice Sheet (figure 1).

ka b. 2000 AD 11.0

12.0

13.0

14.0

15.0

-44-42-40-38-36-34 NGRIP delta18O (per mille)

Fig. 1: δ18O curve for the Lateglacial derived from the NGRIP ice core (based on data from NGRIP members, 2004, Rasmussen et al., 2005, Rasmussen et al., 2006b).

Changes in climate do not always occur simultaneously, ice core records are used to compare timing in marine and terrestrial records with. In terrestrial records, climate changes are recorded by changes in physical and chemical properties of the sediment. Lake sediments are useful for climate reconstruction because they provide a continuous record of changes in ecosystems, lake sediment processes and climate. The environmental change recorded in Ireland’s lake sediments is highly similar to a maritime environment as the climate in Ireland is for a large part affected by the Atlantic Ocean. To explore the correlation between marine, ice core and terrestrial records near western Ireland, this study examines the Lateglacial climate change and chronology in western Ireland 5 and compares the results with marine and ice core records in its vicinity. Several climate proxies were examined to present a reliable estimate of Lateglacial climate at the study site. The climate proxies investigated are organic matter content and isotope ratios (δ18O and δ13C); previous research, ice core and marine records are used for correlation. In addition to this, microtephra analysis was performed to construct a chronological framework.

The following research questions were formulated: 1. How did climate change during the Lateglacial at the study site?

2. Is tephra deposited at the study site and if so, can shards be appointed to a known tephra?

3. How do the results of the multiproxy analysis correlate to previous research in the region?

4. How do the results from this study correlate to results from marine and ice cores?

The thesis is structured as follows. First, Lateglacial climate in northwest Europe and Ireland is presented as known from literature (chapter 2). Next, the study area is described (chapter 3). Chapter 4 elaborates on the methods used. Subsequently, the results of the study are presented and discussed. To end with the conclusion in which the research questions are answered.

As a preparation on this study a literature review was written in a joint effort with Arthur Lutz, some parts of this literature review are literally copied in this thesis. Therefore, chapter 2 is almost completely written by Arthur Lutz, while the rest of this thesis is written by the author.

2. Lateglacial climate

The Lateglacial period is defined as the period in the Quaternary that forms the transition between the end of the last glacial stage, the Weichselien, and the present interglacial stage, the Holocene. The period lasts from ca. 14.0 untill 9.0 ka 14C years BP. The period is characterised by a series of relative rapid cooling and warming events. Figure 2 shows the position of the Lateglacial in the past 150,000 years. The Lateglacial climate can be inferred from ice core records, marine records and terrestrial records.

Fig. 2: The last 150.000 years in the Quaternary period. The Lateglacial forms the transition from the Late Weichselian to the Holocene. (Lowe & Walker, 1997)

2.1 Ice cores From ice cores, an annual δ18O record can be retrieved. Figure 3 shows the locations of past drilling sites in Greenland. The most important are the drilling sites located near the center of the ice sheet, where the best ice cores can be retrieved. The δ18O record from the NGRIP ice core is used in this thesis as a reference to which the later discussed marine and terrestrial records are correlated. The NGRIP δ18O record for the Lateglacial period is shown in figure 4.

Fig. 3: Map of Greenland showing sites where ice cores were retrieved (www.glaciology.gfy.ku.dk).

The INTIMATE (Integration of Ice Core, Marine and Terrestrial Records) group suggested an event stratigraphy for the Lateglacial in the North Atlantic region based on the Greenland ice core records (Björck et al., 1998). The most recently updated version of this event stratigraphy (GICC05 chronology (Lowe et al., 2008)) is shown in figure 4. The period is divided into events and subdivided into episodes, which are all labeled in the figure. Table 1 shows the dates for the different events and episodes. This event stratigraphy forms a framework to correlate ice core, marine and terrestrial records.

Onset Holocene

Fig. 4: δ18O curve for the Lateglacial derived from the NGRIP ice core (based on data from NGRIP members, 2004, Rasmussen et al., 2005, Rasmussen et al., 2006). GICC05 chronology events are indicated (Lowe et al., 2008).

Events Episodes Episodes Ice core age (GICC05) (Continental NW (yrs b2k) for the Europe) onset

Preboreal Oscillation, onset of Holocene Holocene 11.703 Greenland Stadial 1 (GS-1) Younger Dryas 12.896 Pre Younger Dryas GI-1a Warming 13.099 GI-1b Gerzensee oscillation 13.311 Greenland Interstadial 1 (GI-1) GI-1c Allerød 13.954 GI-1d Aegelsee oscillation 14.075 GI-1e Bølling 14.692

Table 1: The GICC05 chronology for key climatic events during the Lateglacial period (After Lowe et al., 2008).

At the end of the Weichselien glacial period temperatures suddenly rise between 14.7 and 14.1 ka b2k. This warm period is the warmest period in the Lateglacial; the Bølling interstadial (GI-1e). From the end of the Bølling interstadial a general cooling trend can be observed until 11.7 ka b2k; the onset of the Holocene. This period with an overall cooling trend is also characterized by relative warm and cool periods. A colder period; the Aegelsee oscillation (GI-1d), starting around 14.1 ka b2k is followed by a warmer period; the Allerød insterstadial (GI-1c), starting around 13.95 ka b2k. The Allerød insterstadial shows some minor oscillations. The period between 13.3 and 13.1 ka b2k is a cold period referred to as the Gerzensee oscillation (GI-1b), followed by a warmer period lasting until 12.9 ka b2k; the Pre Younger Dryas Warming (GI-1a). The coldest part of the Lateglacial period is the Younger Dryas stadial (GS-1), lasting until the onset of the warm Holocene. The Younger Dryas stadial also shows some minor oscillations.

2.2 Marine cores Ice core records and marine records can often be correlated. In this chapter two examples of marine cores retrieved near Ireland will be explained. A 30 m long marine core (MD95-2006) was retrieved from the Barra Fan north-northwest of Ireland (Knutz et al., 2001) (see figure 5 for location). The core extends back to 45 kyr BP. The Barra fan is constructed from debris flow lobes and glacimarine sediments fed by ice streams during the Pleistocene glaciations. Figure 6 shows the results of the core analysis. The results for the Lateglacial period are compared to the NGRIP δ18O record. The same general trend with an initial warming around 15 ka BP followed by a cooling trend lasting until 11.7 ka BP at the onset of the Holocene. The Younger Dryas stadial (GS-1) is clearly observable. The Bølling (GI-1e), Aegelsee oscillation (GI-1d) and Allerød (GI-1c) cannot be separated, but are visible as one warm period (referred to as the Bølling-Allerød interstadial).

Fig. 5: Location of marine cores MD95-2006 and NA 87-22 (Google Maps)

Fig. 6: Correlation of grain parameters and planktonic foraminifera from marine core at Barra Fan to NGRIP ice core oxygen isotope record. (Knutz et al., 2001)

Another core was retrieved off the Irish coast (Duplessy et al., 1996) (see figure 5 for location). For this core, the summer sea surface temperature was calculated from oxygen isotopes in planktonic foraminifera, as can be seen in figure 7. Comparing the sea surface temperature results to the NGRIP δ18O record shows a good correlation. The warm Bølling period (GI-1e) is recognizable as well as the cooling trend until the onset of the Holocene. As with the other marine core, this core is sampled at a relative low resolution, therefore no minor oscillations can be recognized.

Fig. 7: Sea surface temperature and salinity estimates for core NA 87-22 off Ireland (Duplessy et al., 1996). Results are compared to NGRIP δ18O record.

2.3 The Lateglacial in northwest-Europe

Nomenclature The nomenclature and definition of the Lateglacial period differs across Europe (Fig. 8). Except for differences in nomenclature, it is also clear from figure 8 that the period defined as the Lateglacial (Windermere) Interstadial on the British Isles is differentiated further into the Bølling Interstadial (GI-1e), Aegelsee oscillation (GI-1d), Allerød Interstadial (GI-1c) and Younger Dryas Stadial (GS-1) in Continental Northwest Europe. Furthermore, the Lateglacial Interstadial is named the Woodgrange Interstadial in Ireland. The continental North-west European nomenclature will be used in this thesis. However, in some figures the original regional nomenclature is maintained.

Fig. 8: Nomenclature differences for the Lateglacial period on the British Isles and in Continental North-west Europe (After Lowe & Walker, 1997).

Climate Dating of the boundaries between the different climatic periods in the Lateglacial has traditionally been based on 14C dates from bulk samples and Accelerator Mass Spectrometry (AMS) datings from macro- and microfossils (Walker, 1995). A difficulty in 14C dating for Lateglacial material comes from temporal variations in 14C production resulting in ‘plateaux’ in the calibration curve. This results in a high uncertainty for 14C dates at 12.7 ka, 10.0 ka and 9.5 ka 14C years BP (Ammann and Lotter, 1989). The period 14.0-9.0 ka BP is traditionally divided into chronostratigraphic units (chronozones) (Mangerud et al., 1974). Difficulty in the division of the Lateglacial into chronozones comes from difficulties concerning 14C dating, and ambiguity in the chronozones themselves. Chronozones in Europe are defined by radiocarbon dated biozones, which have invariably time-transgressive 14 boundaries and thus vary across Europe. Still, it is possible to use a broad framework for the Lateglacial in northwest Europe (Lowe & Gray, 1980) in which four clearly defined climatic episodes can be recognized: - An initial cold phase prior to ca. 13.0 14C ka BP in which the first signs of warming are apparent (Pleniglacial/GS-2) - A period of significantly warmer conditions from around 13.0-12.5 14C ka BP to ca. 11.0 14C ka BP (Bølling/Allerød Interstadial (GI-1)) - A second cold phase from around 11.0-10.0 14C ka BP (Younger Dryas or Late Dryas Stadial (GS-1)) - Climatic warming of the early Holocene

These climatic episodes are comparable to the climatic signal indicated by the NGRIP oxygen isotope record.

Figure 9 shows the schematic climatic variations in Europe during the Lateglacial. The general picture, indicating a climate warming around 13.0 14C ka BP, followed by a cooling trend lasting until around 14C 10.25 ka BP and climate warming from that point, is comparable for every region in NW-Europe. The regions that are located further away from the Atlantic Ocean show a more pronounced signal compared to the regions located near the Atlantic Ocean. This is especially visible during the Lateglacial Interstadial. In continental Europe this period can be divided into the Bølling interstadial (GI-1e), Aegelsee oscillation (GI-1d) and Allerød interstadial (GI-1c), whereas these fluctuations are less pronounced in regions near the Atlantic Ocean. In those regions this division is difficult. Renssen et al. (2002) combined multiproxy data from different sites in Europe with a numerical model to determine January and July temperatures across Europe during the Late-glacial. During cold phases (Late Pleniglacial (GS-2) and early Bølling (GI-1e)) temperature gradients were very steep in north-south direction, ranging in January from -25 ºC in northern Europe to 5 ºC at the Mediterranean. The temperatures in East-West direction ranged between -15 ºC and -25 ºC. During warm phases (Bølling (GI-1e) and Preboreal) temperature regimes were like the present situation, although temperature gradients were somewhat steeper. Temperature increases during transitions to warmer climate were highest in the northern and western parts of Europe, influenced by latitudinal movements of the North Atlantic polar front and sea-ice margin.

ossil coleopteran curvesossil coleopteran records. The Europe. The curves for Britain and the Netherlands are based on f

Fig. 9: Temperature curves for different regions in northwest for Ireland and northwest Germany are based on palaeobotanical al., data. (Walker et 1994)

Figure 10 shows (a) the Lateglacial and earliest Holocene chironomid inferred mean July air temperature reconstruction (ºC) at Whitrig Bog, southeast Scotland (Brooks & Birks, 2000b), and (b) the NGRIP oxygen isotope record. Mean July air temperature increases from around 6 ºC at the beginning of the Interstadial to a maximum of about 12 ºC. There is a downward trend during the Interglacial with four oscillations of varying intensity, which can be correlated to the NGRIP ice core. The events are labeled as in the GICC05 chronology. Then temperatures rose to about 11 ºC before dropping to about 7,5 ºC at the beginning of the Younger Dryas (GS-1). During the Younger Dryas (GS-1), temperature gradually increased to about 9 ºC before rising rapidly at the onset of the Holocene. As indicated in the figure, the records can be very good correlated. Not only general trends are comparable, also minor oscillations visible in the NGRIP record are distinct in the Whitrig Bog record. Figure 11 shows the chironomid inferred mean July air temperatures (°C) for a record from Hawes Water, northwest England (Bedford et al., 2004). The Hawes water record shows the same picture as the Whitrig Bog record, but indicates a maximum temperature of 13.2 ºC at the start of the Interstadial, also dropping to 7,5 ºC during the Younger Dryas (GS-1). The maximum temperature in the Younger Dryas (GS-1) was 10 ºC according to the Hawes Water record. Like the Whitrig Bog record, the Hawes Water temperature record can be very well correlated to the NGRIP ice core. Not only general trends can are comparable, also minor oscillations visible in the NGRIP record are clearly recognizable in the Hawes Water record.

Fig. 10: Chironomid-inferred mean July air temperature record for Whitrig Bog, southeast Scotland compared to NGRIP ice core oxygen isotope record. (Brooks & Birks, 2000b)

Fig 11: Chironomid-inferred mean July air temperature record for Hawes Water, northwest England (A) compared to oxygen isotope record from the same core (B) and NGRIP oxygen isotope record (C). (Bedford et al., 2004)

Vegetation Figure 12 shows the vegetation development during the Lateglacial for different regions in northwest Europe. At the end of the Weichselien all regions in northwest Europe were mainly vegetated by open ground and tundra communities. At the beginning of the Lateglacial Interstadial this vegetation was replaced by (open) wood and scrub communities in all regions except for Ireland and northwest Germany, where the vegetation was dominated only by scrub communities. Towards the end of the Lateglacial Interstadial vegetation changed in most regions to woodland. The vegetation in Ireland was dominated by open grassland. During the Younger Dryas (GS-1) vegetation in Ireland and Britain changed to tundra and low alpine scrub, while the vegetation in continental northwest Europe was dominated by boreal wood and heath land. At the onset of the Holocene the vegetation changed to woodland in all regions in northwest Europe.

Geomorphology and soils Figure 13 shows a schematic summary of the development of geomorphology and soils during the Lateglacial for different regions in northwest Europe. The main difference is caused by the fact that the British Isles were glaciated during the Late Weichselien and the Younger Dryas (GS- 1), while continental northwest Europe was not. During the Lateglacial Interstadial the landscape became stable and soil development took place in most parts of northwest Europe. Sporadic/discontinuous permafrost seems to have occurred in Ireland and Britain during this period, while there was no permafrost in continental northwest Europe. During the Younger Dryas (GS-1) permafrost occurred in all regions and mineral inwash into lakes took place. Ireland and Britain were partly glaciated. When the Holocene began, soil stabilization and development began in all regions in northwest Europe.

Fig. 12: Vegetation development for different regions in northwest duringEurope the Lateglacial. (Walker et al., 1994)

: Development of geomorphology and soils during the Lateglacial in northwest Europe. (Walker et al., 1994) g. 13 Fi

2.4 The Lateglacial in Ireland

Figure 14 shows the glaciation history for the British Isles during the Weichselien and the Lateglacial period. The Late Devensian ice sheet has disappeared completely from Ireland at 13.0 14C ka BP (Lowe & Walker, 1997). The sites in Ireland referred to in this chapter are indicated in figure 14. Research results consist of pollen diagrams, lithology descriptions, organic content analysis, whole core volume susceptibility, chemical element concentrations, macrofossils (AMS dates), mollusk and osctracod analysis, stable isotope records and carbonate content.

Chronology In Ireland, 14C datings for the end of the Lateglacial Interstadial (Woodgrange Interstadial) show intersite variations. Dates spanning the interval from ca. 10.6 to 11 ka BP are found for the onset of the Younger Dryas (GS-1) (Craig, 1978; Cwynar & Watts, 1989). Dates for the end of the Younger Dryas also vary. Cwynar & Watts (1989) found dates in the range of 9.5 to 9.9 ka BP and dates up to 10.2 ka BP have been reported by Craig (1978). The Lateglacial Interstadial (Woodgrange Interstadial) is not often divided into Bølling (GI-1e) and Allerød (GI-1d) periods because of confusions over the climatostratigraphical and chronostratigraphical application of these terms (Walker et al., 1994).

Fig. 14: Glaciation limits during recent glaciations in Britain. (Bowen et al., 2002). Sites referred to in the text are indicated in the figure.

Climate The temperature curve for Ireland shown in figure 9 is based on palaeobotanical data (Walker et al., 1994). The curve shows a climatic warming around 13 ka BP, with maximum summer temperatures of at least 14-15 ºC, comparable with present summer temperatures in southern Ireland. The following cooling in climate seems to have a step-like pattern. During the Younger Dryas stadial (GS-1) mean annual air temperatures fell below 10 ºC, and January mean temperatures of -20 ºC have been inferred for the British Isles (Atkinson et al., 1987). A rapid rise in temperature at the onset of the Holocene is reflected in all available forms of proxy data. Ahlberg et al. (1996) studied oxygen isotopes for Red Bog and Lough Gur in western Ireland. Figure 15 and 16 show the studies’ results correlated to the NGRIP δ18O record. The δ18O record from Red Bog and Lough Gur gives an indication for the relative temperature changes during the Lateglacial. Only relative temperatures can be estimated and the size of the error margin is uncertain. The periods Oldest Dryas, Bølling (GI-1e), Aegelsee oscillation (GI-1d), Allerød (GI-1c), Younger Dryas (GS-1), Preboreal and Boreal are clearly recognizable. The Gerzensee oscillation (GI-1b) is also visible in both records. The correlation with the NGRIP δ18O record is good for both sites.

Fig. 15: LOI and Oxygen isotope record for Red Bog compared to the oxygen isotope record from NGRIP ice core. (After Ahlberg et al., 1996)

Fig. 16: Oxygen isotope record for Lough Gur compared to the oxygen isotope record from NGRIP ice core. (After Ahlberg et al., 1996)

A multiproxy analysis was carried out in 1999 on a core from Tory Hill basin in western Ireland (O’Connell et al, 1999). The core was analyzed for pollen, macrofossils, loss-on-ignition, magnetic susceptibility, carbonate content, oxygen and carbon isotopes, and heavy minerals. Environmental change, including estimates of summer temperatures, could be reconstructed. Figure 17 shows a part of the analysis’ results. Ages were assigned to biostratigraphical and lithostratigraphical boundaries by cross-correlation with the Late-glacial event stratigraphy put forward by Björck et al. (1998). The estimates of summer temperatures are based on the δ18O record and the vegetation record (pollen and macrofossils). The period from 14.6 to 12.7 ka BP represents a cold, steppic environment from the last phase of the Pleniglacial (GS-2). A thermal maximum was achieved early in the interstadial with a maximum summer temperature of 14 ºC. Within this warm period, the proxies suggest at least one substantial, short lived climatic 26 oscillation. The later part of the Bølling/Allerød interstadial (GI-1, after ca. 12 ka BP) began with a temperature decrease, maintained until just before the Younger Dryas (GS-1) where a temperature rise is recorded in the proxies, referred to as a pre-Younger Dryas warming (GI-1a). During the Younger Dryas stadial (GS-1) summer temperature drops to around 6 ºC. The early part of the Younger Dryas stadial (GS-1) is not well represented by the stable isotope record, for reasons not understood. The period which was referred to as the pre-Younger Dryas warming (GI-1a) is also visible in the other proxy records. In figure 17 the results are correlated to the NGRIP δ18O record. The results from Tory Hill show the same general trend as the NGRIP δ18O record. The Bølling Interstadial (GI-1e), pre Younger Dryas warming (GI-1a), and Younger Dryas (GS-1) are clearly recognizable. Minor oscillations are not well recognizable in the Tory Hill records. The researchers found that the variations in δ18O did not directly reflect temperature changes at the study site, but rather δ18O variation in precipitation, which was forced by temperature and changes in atmospheric circulation.

Fig.17: Percentage pollen curves for Betula and Juniperus, and δ18O, erosion, and July temperature values compared to NGRIP oxygen isotope record. (After O’Connell et al., 1999)

At another lake near the study site climate change is reconstructed. A 7.6 m long core was retrieved from the southeastern shoreline of Lough Inchiquin in 2002 (Diefendorf et al., 2006). The results from the analyzed core are presented in figure 18. Ages were connected to the core by 14C dating of bulk aquatic macrofossils, bulk calcite and organisms. The Lateglacial chronology is only based on two 14C datings. Since the datings are influenced by the hard-water effect, they were corrected for an assumed constant hard-water effect, which was quantified by comparing contemporaneous wood and carbonate macrofossils. According to these results the climate in Ireland was highly variable during the Lateglacial and early Holocene. Some previously unidentified climate anomalies in western Ireland were identified at 10.8 cal yr BP and 7.1 cal yr 18 BP. After the Younger Dryas (GS-1), δ Ocalcite values recover strongly, followed by decreasing values. The most severe climate decline in Ireland during the Younger Dryas (GS-1) appears much later than in other records, which was also noted at Tory Hill by O’Connell et al. (1999). The δ18O record is correlated to the NGRIP δ18O record in figure 18. The overall trend from the NGRIP record is also visible in the record from Lough Inchiquin. The Younger Dryas event (GS-1) is visible as well as the pre Younger Dryas warming (GI-1a). The timing of the events sometimes differs in both records, which could be partially caused by the limited datings on which the Lateglacial chronology at Lough Inchiquin is based, and therefore it is difficult to correlate both records.

18 Fig. 18: Lithology (A), weight percent total calcite (B), weight percent total organic matter (C), δ Ocalcite values vs. age in cal yr BP with significant climate events labeled in brackets and a 50-year moving average in grey inllustrating first order 28

18 trends, and the NGRIP δ Oice record for comparison. Ice-rafted debris is labeled as IRD. High values indicate higher temperatures and low values are interpreted as lower temperatures resulting from Rayleigh distillation/atmospheric 18 temperature relationships that force δ Oprecipitation values. (After Diefendorf et al., 2006)

Limited data is available on former precipitation but it might be inferred that Ireland formed a more oceanic region within northwest Europe with relatively high precipitation (Walker et al., 1994). There are no temperature records based on chironomids for Ireland, but there are chironomid records from northwest England and southeast Scotland (see chapter 2.3).

Vegetation Palaeobotanical records from Ireland show that the largest part of the Lateglacial vegetation was characterized by herbaceous and low shrub vegetation with very few trees (Watts, 1985). The late Weichselian tundra vegetation was first succeeded by a Rumex-Salix vegetation, then around 12.4 ka BP by scrub and heath land and finally by open grassland. Pioneer trees were never abundant. Pollen assemblages from Tory Hill show a distinct rise in Betula and Juniperus during the Bølling Interstadial, indicating this was the only period during the Lateglacial with substantial development of woody vegetation (O’Connell et al., 1999). Apart from climatic reasons, tree growth might also be tempered by grazing pressures from large animals like the giant Irish deer and reindeer (Mitchell, 1986). At some sites in Ireland and Scotland the mid- Interglacial pollen records show a decline in Juniperus and rise of open-ground communities (Watts, 1985; Walker & Lowe, 1990). This characteristic is not as distinct as it is in records from continental northwest Europe (Andrieu et al., 1993). The Younger Dryas in pollen records from Britain and Ireland is characterized by the replacement of wood and heathland taxa by tundra and low alpine scrub communities (Watts, 1985). The transition from Younger Dryas to Holocene is characterized by a succession from open ground communities to closed woodland (Walker et al., 1994,).

Geomorphology, soils and lithology Lateglacial lake sediments in Ireland have a distinct stratigraphy separated into three parts. The main lithological features are (O’Connell et al., 1999): - Silt/clay sediments from the final phase of the Pleniglacial (GS-2) followed by - Organic-rich gyttja or marl, depending mainly on the local bedrock, and - Clay/silt sediments during the Younger Dryas stadial, ending with a sharp transition to - Organic-rich gyttja or marl at the beginning of the Holocene.

Between 12.0 and 11.8 ka BP there are indications for renewed minerogenic inwash into lake

29 basins in Ireland. This is particularly true for upland areas (Pennington, 1977). Possibly, discontinuous permafrost occurred in Ireland during the Younger Dryas (Bryant & Carpenter, 1987). Also, the Lateglacial sequences are often thick (usually more than 1 m); sometimes sequences are longer than 2-4 m (Singh, 1970). An example of a Lateglacial sediment sequence at Tory Hill is shown in table 2 (O’Connell, 1999).

Depth (m) Sediment type

0.00 – 1.00 Black fibrous fen peat

Holoc. 1.00 – 3.70 Lake marl Sediment

3.70 – 6.00 Blue/grey clay and silty clay (Younger Dryas) Grey calcareous clay (Younger 6.00 – 6.20 Dryas/Interstadial transition) Lake marl with clay (8.52–8.96 m (end of Bølling, cf Older Dryas) and 9.48-9.76 (within 6.20 – 10.04 6.20 - 9.86 (Predominantly Bølling) and organic enriched (7.88-8.08, 8.35- lake marl; 8.52 and 8.96-9.48) layers Interstadial) Grey calcareous clay, organic rich; much 9.86 – 10.04 clayey silt between 10.00 – 10.04 Dark blue clay with faint irregular veins, 10.04 – 10.12 strong H2S smell Blue clay with fine laminations showing grey blue/black couplets when the core was first

Lateglacial sediment sediment Lateglacial 10.04 – 13.52 10.12 – 10.36 sliced open. On oxidation, the couplets were (silty clay) no longer visible (Pre-interstadial first recorded in spectrum 1008 cm) Pink clay, with fine laminations showing 10.36 – 13.52 couplets consisting of a grey sandy silt layer overlain by a brown clay layer 13.52 – 13.60 Grey gravelly limestone sand

Table 2: Stratigraphy Tory Hill (O’Connell et al., 1999)

3. Study area

The study area (grid ref. R 41785 96437; 8º 52’03.0” W, 53º 00’55.8” N; EPE 4), is situated near the west coast of Ireland in the County Galway, close to the County Clare boundary; and is called Lurga after the nearby village. More precisely, it is located north of Tubber, in the south-eastern Burren (figure 19 and appendix A) and is part of a shallow basin situated in a karstic limestone area. The shallow north-south running basin is partly filled with marl overlain by peat and is a cutover bog. In several places, the marl deposits include Lateglacial sediments. Because the site is still partially a lake, a part of the area is not accessible. In previous research (Paus et al., 1994) a Lateglacial sediment record was found at the study site on which already some research was done. The sediment record at Lurga seemed suitable for a Lateglacial climate reconstruction of western Ireland and the construction of a chronological framework.

Study Study site site

Fig. 19.A Map or Ireland with the study site indicated. Fig. 19.B: Topographical map of study site Lurga (Ordnance (www.climate-zone.com) Survey Ireland)

Fig. 19.C: Detailed map of the study site Lurga. The black solid line indicates the transect at the study site and the studied sediment core is indicated with LU-A.

4. Methods

Several transects and individual corings were made to explore the subsurface of the area (appendix 1 and fig. 19.C). At the location where the lake infill was the thickest and a Lateglacial sediment sequence seemed to be present, a 6-cm diameter core (LU-A) from 611 to 764 cm under the present surface was retrieved for further laboratory analysis. As the coring was done with an adapted (Livingstone) piston corer (which can retrieve a sediment sequence of one meter in length), an overlapping core was taken to replace the disturbed top and bottom parts of the one meter long sediment sequences. In the laboratory the following investigations were carried out. A detailed lithostratigraphical description was made; the organic matter content, carbonate concentrations and oxygen- and carbon isotope ratios were measured; and finally potential tephra shards were extracted for optical identification.

4.1 Climate reconstruction

4.1.1 Lithostratigraphy The lithostatigraphy of the sediment core was described accurately (up to half centimeters). The studied sediment core was comprised out of lake sediments. Lake sediment sequences offer a detailed record of environmental change. Given sufficient time, all lakes become infilled with sediments to form mires and bogs. Lake sediments are partly derived from the organic production within the lake and partly from the organic and clastic material washed in from the catchment. The variations in the clastic material in the lake sequence reflect the changing processes acting upon the slopes around the lake catchment. During cold periods more material will be eroded and the relative amount of clastic material in a lake will be higher and of a larger grainsize (Lowe & Walker, 1997). If the subsoil surrounding a lake consists of limestone, carbonate can be dissolved by the groundwater and precipitated in the lake through CO2 uptake of aquatic plants. The resulting deposit is called a calcareous gyttja or a lake marl (Bohncke and Hoek, 2007).

4.1.2 Loss on ignition To asses the biological productivity during deposition, the organic matter content was determined. Loss on ignition (LOI) is a widely used method to estimate the organic matter content of the sediments. The following series of procedures was followed. First, contiguously 1 cm3 samples were taken from the sediment core and placed in weighted crucibles. In order to estimate the organic matter content and not include the water content, the samples were oven-dried to constant weight (12-24 hours at 105 °C) (Heiri et al., 2001). When the samples were dried, they

33 were weighted at a temperature of 105 °C and afterwards placed in a furnace (Heraeus Hanau KR 170 furnace). In the furnace organic matter was combusted to ash and carbon dioxide at a temperature of 550 °C for four hours and subsequently weighted. The LOI value was calculated with the following equation:

LOI550 = ((DW105 – DW550)/(DW105 - Wc)) * 100 %

Where LOI550 is the LOI at 550 °C (as a percentage), DW105 represents the dry weight of the sample before combustion, DW550 represents the dry weight of the sample after heating to 550 °C and Wc is the weight of a crucible (all in grams). The weight loss is in this way proportional to the organic matter content in the sample (Heiri et al., 2001).

Although LOI is considered to be a simple method to estimate organic matter content, factors like exposure time, sample size, position in the furnace and mineral dewatering affect the LOI result (Heiri et al., 2001). These factors were accounted for as much as possible. The exposure time of two hours commonly used for heating to 550 °C may not be sufficient for sediments with high organic matter content ánd high carbonate content; in that case an exposure time of four hours would be more appropriate (Heiri et al., 2001). As the lake infill comprised for a large part out of calcareous gyttja, the carbonate content was very high. The organic matter content was thought to be fairly low, indicating that both conditions, high carbonate ánd organic matter content, are not filled. Nevertheless, to be positive that the low exposure time of two hours was not too short for the complete combustion of the organic matter content, the exposure time was increased to four hours. Also the position of the samples in the furnace influences the weight loss, which is different per furnace. Samples at the centre of the furnace have a higher LOI than samples on the side, which becomes increasingly important with longer exposure times (Heiri et al., 2001). The extent of this effect was assessed in the study. The final factor influencing the LOI result is that, while the organic matter is combusted to ash and carbon dioxide, other reactions can take place which can lead to a weight loss, e.g. dehydration of clay minerals or metal oxides, loss of volatile salts, or loss of inorganic carbon in minerals such as siderite. The mineral dewatering component can be decreased by combustion at lower temperatures; nevertheless the organic matter will not be combusted completely at these temperatures. Therefore the LOI will be a compromise between mineral dewatering and organic matter combustion (Boyle, 2003). In this study the temperature of combustion was not lowered to correct for mineral dewatering.

4.1.3 Isotope analysis Oxygen and carbon isotopes of bulk carbonates were analysed. Isotope analysis on oxygen and carbon isotopes was done to asses past environmental and climate change in the study area. However, there are many non-climate or –environmental factors which can influence the ratios of the isotopes.

Carbonate content The isotope analyses were performed on carbonates, therefore the carbonate content was a boundary condition for isotope analysis and was determined using the Scheibler method.

The Scheibler method is based on the volumetric measurement of CO2 emerging from the reaction of carbonates and HCL, assuming that all the CO2 emerges from the disintegration of

CaCO3 and not from other carbonates (ÖNORM L1084).

+ 2+ CaCO3 + 2H → Ca + H 2O + CO2

The Scheibler method used consisted out of the following steps. 1. A sediment core was sampled continuously with ca. 1 cm3 large samples which were placed in test-tubes. 2. In order to eliminate moisture, the samples were oven-dried to constant weight 3. The samples were grinded to homogenize the sample and to increase the surface area of the sample, to stimulate chemical reaction speed. 4. Samples of 0,2 grams each were placed in Erlenmeyer flasks and diluted with ≤ 15 ml distilled water. 5. Isolated in small cans, the HCl was placed in the Erlenmeyer flasks. 6. The Erlenmeyer flasks were connected to the Scheibler device and the HCl was mixed with the samples, this causes a reaction between carbonates present in the sample and the HCl.

7. The emerging CO2 was measured volumetrically.

8. Measurements on a standard NaO3 -solution were done throughout the course of the day to eliminate any influence of atmospheric pressure differences.

9. The amount of emerged CO2 was calibrated to carbonate per volume sample. 10. The results were corrected for the carbonate content of the used distilled water.

Carbonate can be present in a lake sediment as allogenic, authigenic and biogenic carbonate. Allogenic carbonate is detritus derived from the terrestrial environment washed in from the shores of the lake. Authigenic carbonate is precipitated on the bottom of the lake by the utilization of carbon dioxide for photosynthesis and the resulting supersaturation of calcium carbonate in the

35 water column; it typically forms during summer photosynthesis when biological activity is highest. Biogenic carbonate has accumulated in organisms living in the lake (Leng and Marshall, 2004). It is often assumed, when dealing with a sediment core containing authigenic carbonate, that all the carbonate present is authigenic; this condition is not always fulfilled. Commonly the allogenic carbonate has a larger grain size and is easy to distinguish between the authigenic carbonate. In Lateglacial sediments from Tory Hill (western Ireland), it seems that in a part of the core allogenic material is present. This theory is supported by the poor correlation between δ18O and other proxies where the supposed allogenic carbonate is present (O’Connell et al., 1999). Furthermore, carbonate rich sediment can contain variable proportions of biogenic and authigenic phases with very different compositions, which can hamper the interpretation of the isotope signals. Finally, many palaeoclimatic studies have assumed that authigenic carbonate formed in isotopic equilibrium but few studies have demonstrated that this was the case (Leng and Marshall, 2004).

Oxygen isotope ratios Oxygen has three stable isotopes that occur naturally on Earth: 16O, 17O and 18O. 16O is the lightest and most abundant (99.8%) isotope, 18O is the heaviest isotope and accounts for most of the remaining 0.2%, 17O makes up only a fraction of the oxygen on earth and will not be further discussed. The ratio of 18O and 16O is about 1/400 (Ruddiman, 2001). Small variations around the average ratio of 18O and 16O were measured from lacustrine calcite.

Dried and grinded carbonate samples were exposed to acid and the emerging CO2 gas was analyzed with a mass spectrometer (Siegenthaler & Eicher, 1986). Samples were taken every centimeter. Measurements of the 18O/16O ratio were reported as parts per thousand (‰) from a laboratory standard (Ruddiman, 2001):

18 16 18 16 ()()O / O sample − O / O s tan dard 18 δ O = 18 16 ×1000 ()O / O s tan dard Typical δ18O values of lake sediments fall in the range of -10‰ to -5‰ (Drummond et al., 1995).

General link between δ18O and temperature In the atmosphere there is a gradual poleward transport of water vapour. The isotopic composition of the water vapour will change because of the combination of temperature and the weight difference of the 18O and 16O isotopes. First, because the 18O isotope is heavier, it evaporates less readily from the ocean, leaving the water vapour in the air depleted in 18O. Furthermore, when the air becomes cooler it cannot hold as much water as under warmer conditions and just as the heavier isotope evaporates less readily, it condensates more easily (fractionation); leaving the air even more depleted in 18O. During glacial periods temperatures are

36 lower and the air can hold much less water vapour, resulting in 16O enriched precipitation in the high latitudes. Because the temperature influences the ratio of 18O and 16O, the δ18O signal from authigenic carbonate can act as a proxy for temperature (Ruddiman, 2001). Therefore δ18O is used in many studies as an indicator of palaeotemperature at the study site, but this can only be assumed if a wide range of interlinked environmental processes which can change this ratio of the lake water are negligible or can be corrected for (Leng and Marshall, 2004).

Factors influencing the δ18O of lake water

1. δ18O of Precipitation When there is climate fluctuation the δ18O of precipitation will change as a result of fractionation of oxygen isotopes, but this may be independent of changes in temperature at the study site. For example, if the temperature has risen at the ocean where the source water is evaporated, but the temperature at the study site has not changed, δ18O values will indicate an increase in temperature at the study site while this is not the case. Similarly, a relative depletion of 18O in the sea water will result in an increase in δ18O, independent of temperature changes at the study site. Furthermore, atmospheric circulation may change and increase the distance the water vapour has to travel from the source to the lake. Through increasing this distance, there is more time for the relative heavy 18O isotope to condensate which will indicate a cooling at the study site while this is not the case (Leng and Marshall, 2004).

2. CO2 exchange between the atmosphere and lake water 18 δ O can only be used as a proxy for temperature if the CO2 exchange used for carbonate precipitation in the lake is in isotopic equilibrium with the atmosphere. In many mid-latitude lakes the δ18O of the precipitated carbonate will be mainly influenced by temperature and will only be 18 dampened by the disequilibrium of CO2 exchange; the δ O of carbonates for these temperate latitude lakes will co-vary with temperature with an increase of ~0.36‰ / °C (Leng and Marshall, 2004).

3. δ18O enrichment through evaporation In closed lakes where water is lost mainly through evaporation, water has a variable and elevated oxygen isotope composition. Lake δ18O values are always higher than those of local precipitation as the lighter isotope 16O is preferentially lost to evaporation. In closed lakes, the isotope values are highly variable as the ratio between evaporation and precipitation changes with climate. It is therefore important that an oxygen isotope record, studied to reconstruct palaeotemperature, is not from a closed lake. To determine the degree of closure of a lake through time the correlation

37 between δ18O and δ13C is used; although this generally only applies for lakes that have been closed for long periods (Leng and Marshall, 2004).

4. Basin versus catchment size In small, short residence time lakes seasonal variation in the δ18O values of precipitation is likely 18 to be of large importance for the δ Ocarbonate data, as these small lakes often have highly fluctuating δ18O values. These fluctuations are caused by the daily and seasonal variation of the δ18O values of precipitation. At temperate latitudes winter rainfall has significantly more negative δ18O values than summer rainfall. If a lake is very small in relation to its catchment (with residence times of < 1 year) it is not big enough and well enough mixed to ‘average out’ the short term variation in isotopic composition of the precipitation and hence the isotopic composition will not reflect the δ18O of the mean annual precipitation. Even in lakes with relatively long residence times isotope compositions can reflect summer rainfall rather than mean annual rainfall because of stratification of the water column (Leng and Marshall, 2004).

Carbon isotope ratios On Earth there are three naturally occurring carbon isotopes: 12C, 13C and 14C. The lightest isotope 12C is most abundant in nature (99%), 13C represents 1% while the heaviest and more scarcer isotope 14C is radioactive and is used for 14C-dating, 14C will be further discussed in chapter 4.2.1. The average value of the 13C/12C ratio is less than 0.01 (Ruddiman, 2001). Changes around the average value of the ratio of 13C and 12C were measured in the same manner as δ18O. Similar as with the oxygen isotopes, the 13C/12C ratio is reported as parts per thousand (‰) from a laboratory standard (Ruddiman, 2001):

13 12 13 12 ()()C / C sample − C / C s tan dard 13 δ C = 13 12 ×1000 ()C / C s tan dard δ13C values of lake sediments are in the range of -12‰ to 1‰ (Drummond et al., 1995).

Inorganic δ13C in lake waters reflects environmental change, which is often related to climate change. Inorganic carbon is defined as, carbon which is not produced by an organism, but by inorganic sources such as carbonate mineral precipitation (lake marl). Inorganic carbon in lake water results from the interaction of groundwater with rocks and soils in the catchment. The inorganic carbon precipitates in the lake as authigenic carbonate. The δ13C of authigenic carbonate has almost the same values as the inorganic carbon (there is a negligible difference through isotope fractionation). During this study bulk carbonates are measured therefore biogenic and allogenic can also be present in the samples.

There are two main processes which control the δ13C of inorganic carbon in limestone areas (Leng and Marshall, 2004): 1. The isotopic composition of lake water 2. Photosynthesis/respiration of plants in a lake

Factors influencing the δ13C

1. The isotope composition of lake water There are different types of inflowing waters which can influence the δ13C of the total dissolved inorganic carbon (TDIC) in the lakewater, for instance groundwater and rivers. Such inflowing waters have relatively low δ13C values between -10‰ and -15‰. Soil development can have a large effect on δ13C values of the TDIC because different plant species can have quite different δ13C values (Romanek et al, 1992). In karstic limestone areas δ13C values can be much higher because a proportion δ13C originates from the dissolution of limestone (Andrews et al., 1997), up to 3‰ (Hudson, 1977). At Lough Inchiquin (western Ireland) the catchment changed from exposed limestone bedrock to a vegetated surface, which changed the amount of weathered carbon entering the lake dramatically (Diefendorf et al., 2008). As the study area is situated in a limestone area, high δ13C values were expected.

2. Photosynthesis/respiration of plants in a lake The main controlling factor on fluctuations of the δ13C in the lake water is normally assumed to be the biological productivity, which is highly dependent on temperature. However, aquatic plants can also effect the δ13C of the lake water (Leng and Marshall, 2004). Firstly, aquatic plants often preferentially take up 12C, leading to a depletion of 12C in the water and therefore to higher δ13C in the precipitated carbonate. Secondly, plant photosynthesis often flourishes at summer causing an enrichment of 13C during summer, which results in a seasonal fluctuation in the δ13C. Finally, if stratification in the water column becomes permanent, it can cause a large change in the δ13C values of precipitating carbonate. Stratification can reduce the rate of organic matter oxidation causing higher δ13C values (Leng and Marshall, 2004).

4.2 Chronological Framework

Because climate changes were abrupt during the Lateglacial, the effective site-to-site correlation should be at least with a decadal precision (Davies et al., 2002). Radiocarbon dating does not provide this precision because errors arise through site specific factors (e.g. hard-water effect in lakes or reservoir ages in marine sequences), sample selection and laboratory measurements (Lowe and Walker, 2000). Furthermore, imprecision may be introduced by the calibration of radiocarbon dates, since there are uncertainties in the dataset (f.i. radiocarbon plateaus) which decrease the precision drastically (Davies et al., 2002). Tephra horizons are time-parallel horizons which can serve as important ‘tie lines’ between sites. Some tephras are already securely, high-precision dated (such chronologies are called regional master sequences) and can be used to rapidly correlate and date new records or test the reliability of calibrated radiocarbon data sets. However, many tephras are not reliably dated yet (Lowe, 2001). Several 14C-plateaus are present during the Lateglacial, making the precise timing of climate fluctuations difficult to asses. The construction of a tephro-chronological framework for Ireland ensures more precise dating of the Lateglacial, furthermore the synchronism of Lateglacial climate shifts could be tested (Turney et al., 2004). To construct a tephrochronological framework potential tephra shards were extracted from the sediment core to be optically identified.

4.2.1 Tephra The originally Greek word ‘tephra’ means ash and is used in earth sciences as a term for all airborne pyroclasts, including both air-fall and pyroclastic flow material. Tephra includes a wide range of different sized fragments, but owing to the long distances from active volcanoes, ash- size fragments are predominantly studied (Haflidason et al., 2000). Because tephra layers are deposited almost instantaneously spatially, they represent time-parallel marker horizons within stratigraphical sequences (Davies et al., 2002). Several factors influence the deposited shard concentration such as distance from the volcanic source, precipitation patterns during the eruption and subsequent tephra dispersal, and sedimentary processes within the catchment (Davies et al., 2002). The last fifty years tephra is mostly studied as visible ash layers, in recent years tephra is discovered and studied at the microscopic level (Hall and Pilcher, 2002). These microscopically small shards, typically smaller than 100 μm, are also referred to as ‘microtephras’ or ‘cryptotephras.’ Tephra is rhyolitic and basaltic in composition; the rhyolitic (lighter) component is normally more abundant (Pyne-O’Donnell et al., 2008). With tephrochronology the rate and synchronism of climate change can be assessed, it can be used as a chronological framework which facilitates high-precision correlations between

40 sites (Pyne-O’Donnell et al., 2008). In recent years, microtephra is discovered more widespread than previously thought. Scientists are still discovering microtephra at more distant locations, which increases the spatial high-precision relationships between sites and increases the potential of tephrochronology (Hall & Pilcher, 2002; Turney et al., 2004; Wåstegard et al., 2006). Individual tephra layers can be assigned to a particular volcanic event on the basis of a combination of criteria, including the stratigraphic position of the layer and independent assessment of the age of the layer as well as geochemical properties and shard morphology (Davies et al., 2002).

Chemical properties and shard morphology Glass shards formed during rapid cooling in a volcanic eruption are thought to have approximately the same composition as the bulk geochemistry of the magma. If a tephra deposit can be shown to have a distinct geochemical fingerprint, it can be used to correlate the stratigraphic units in which it occurs and if this tephra is securely dated the age can be estimated. Tephra deposits in Ireland have principally an Icelandic origin, but in consideration of recent more widespread findings of microtephra, it is possible that tephra originating from the Eifel (e.g. Laacher See tephra) is also deposited. Overall, each ash province has its own distinct geochemical composition; and even between individual volcanoes often a distinction can be made (figure 20) (Davies et al., 2002).

Fig. 20: Bo-plot of SiO2 and K2O concentrations in tephras derived from the main European ash provinces. (Davies et al., 2002)

Not all deposited tephras can be tied to a known volcanic source. The reasons for this can be that they are new tephras or that there is an error in the chronological framework (Haflidason et al., 2000); however they still provide reliable methods of site linkage (Hall and Pilcher, 2002). 41

Another discriminator of tephra deposits is the shard morphology. The shape of tephra shards is difficult to specify, therefore it is described in three categories. The first one is based on axial ratios and includes parameters such as sphericity and elongation. The second category involves the overall roundness or angularity of the shards, and the third is concerned with small- scale surface textures. The grain morphology is mainly controlled by two processes: the expansion of gas in fluid magma, and breakage due to forces acting on the particles during eruption. After initial deposition, tephras may be subjected to erosion, transport and re-deposition just as any other sediment (Haflidason et al., 2000). All these processes modify the grain morphology and can be assessed visually with classification schemes (Heiken and Wohletz, 1985).

Alternation of tephra shards under natural conditions Glass corrosion takes place in natural environments because of different reasons. For example, volcanic glass shards can be altered through biogenic grooving or microbiological activity; this effect is most pronounced on rhyolitic particles (Haflidason et al., 2000). Secondly, the acidity of the sediment plays an important role. In mildly basic to acidic environments (pH<9) the predominant mechanisms of chemical shard alternation is ionic exchange of cations at the glass surface. The lost cations result in the formation of a leached Si gel layer. In more alkaline environments (pH>9) the predominant chemical alternation process is network dissolution (hydroxyl ions in solution disrupt siloxane bonds in the glass surface), which leads ultimately to the dissolution of the glass shards. Finally, the durability of glass shards, occurring in a sediment sequence, depends on the degree in which an inert Si gel layer forms on the surface of the shard. Natural chemical alternation of tephra shards leads to visible alternation of the shards which can hamper the analysis of the shards (figure 21) (Blockley et al., 2005).

Fig. 21: Basaltic tephra shards of the Vedde ash from Loch Ashik, Scotland, exhibiting examples of chemical alternations: (A) the formation of a leached Si gel layer, (B) pitting erosion through cation removal, (C) preferential leakage of particular phases of the glass and recrystalisation of leached products onto the glass surface. (Blockley et al., 2005)

Micro tephra extraction techniques The problem of using microtephra is that they are not visible to the naked eye, microtephra is especially difficult to distinguish when deposited in mineral rich lake sediments (Turney, 1998). Different separation techniques are used to detect and isolate tephra particles from sedimentary sequences. These include ashing of organic matter (Pilcher and Hall, 1992), digestion of biogenic and opaline silica (Rose et al., 1996), the use of magnetic properties of ash layers (Pawse et al., 1998), X-ray analysis (Dugmore and Newton, 1992) and a density separation (flotation) technique (Turney, 1998), which has been modified by Blockley et al. (2005). The latter of these extraction techniques is the most promising (especially in mineral rich sediments) and has therefore been used for this study. With all micro tephra extraction techniques the sample is ‘cleaned’ before tephra shards can be extracted, there is evidence that some ‘cleaning’ methods affect the chemical stability of tephras or lead to degradation of the tephras. For example, during ashing the chemical composition of tephra shards can be altered and this should therefore be avoided if the shards will be chemically analyzed. Another problem with ashing is that clay-rich samples can become solid like bricks, which makes further analyses impossible. Alternatively, samples can be treated with hydrogenperoxide (H2O2) to remove organics (Koren et al., 2008). Furthermore, as earlier described, aggressive acidic and alkaline solutions alter the glass shard’s chemical properties and should therefore be avoided in the preparation of tephra glass shards for (optical and) geochemical analysis. Some scientists claim that only the surfaces of the shards are affected and that polishing exposes the original chemical composition of the glass shards. Blockley et al. (2005) tested this hypothesis with several experiments (figure 22) which all resulted in a significant change in the chemical composition or decrease of the number of shards although the shards were polished.

Fig. 22: Sample loss due to NaOH preparation on 20 paired samples of glass extracted from a bulk sample of the Baia Averno tephra. (Blockley et al., 2005)

Density separation technique The density separation technique purifies tephra without the need for aggressive chemical solutions which can alter the chemistry of tephra shards or dissolve them all together. The technique is based on stepped flotation, with the use of a series of solutions of sodium polytungstate which increase in specific gravity to remove any impurifications. This stepwise flotation method has several advantages. First, concentrations of recovered tephra shards are higher. Second, there is less detrital material present, so that tephra shards are less obscured. Finally, tephras have cleaner surfaces, so that distinguishing surface features is easier. The achievement of this technique increases the confidence with which decisions can be made about whether certain tephras are present in, or absent from, a sediment sequence (Blockley et al., 2005). There are however still some problems concerning the density separation technique (Blockley et al., 2005). 1. The method only reduces the potential for chemical alteration in the laboratory; little can be done about chemical alteration under natural conditions. 2. No liquid medium is chemically inert. Hence, some degree of chemical alteration is inevitable. 3. The specific density of tephras can be altered under natural sedimentary conditions through the formation of, for example, Si gel layers on the surface of shards. Therefore, the optimal specific gravity for the recovery of a particular ash type discovered for one sequence may not apply in another case, where more or less chemical alteration of the shards has taken place. 4. Some phytoliths (opaline Si) look remarkably like tephra shards, and can mimic their behavior under cross-polarized light when Canada balsam is employed as the mounting medium, therefore glycerol was used with which this mimicking behavior does not occur. Although most phytoliths are removed in the final cleaning float, in some instances the specific gravity of phytoliths can be as high as that of tephra. It can thus be difficult to isolate and concentrate some tephras using the density separation method, though chemical digestion procedures are even less efficient at solving this particular problem. 5. During the purification, basaltic tephra is not isolated because its specific gravity is approximately the same as quartz grains, which are removed. With other techniques the basaltic tephra can be isolated from the residue.

Applied tephra extraction and optical identification procedure The procedure of tephra extraction and optical identification of shards is explained below step by step. 1. The organic matter content was removed by ashing the samples at 550 ºC for four 44

hours (the samples used for the loss on ignition method were used). Samples with high clay content became too hard after combustion (‘bricks’) to perform further analysis on. To replace these ‘bricks’, new samples were taken and the organic matter content was removed (after step 2!) by adding 2-4 ml hydrogen peroxide and heating it at 120 ºC until all organic matter and hydrogen peroxide had reacted to

form CO2 and H20. 2. Calcite was removed by adding 20 ml of 5% HCL solution and giving the HCL sufficient time to react with the sample (usually overnight). 3. All the sediment smaller than 31 μm (mainly clay) was removed by sieving. 4. In some samples biogenic silica were present. These were removed by adding 15 ml of 0.5 M NaOH and heating it for 3 hours at 70 ºC. 5. Centrifuge test-tubes with the samples in it were filled with distilled water to 40 ml, stirred and centrifuged at 2500 rpm at brake rate 0 for 20 minutes. Afterwards the distilled water was decanted. 6. The test-tubes were filled with 15 ml sodium polytungstate with a density of 2.5 g/cm3, the sodium polytungstate was mixed with the distilled water by stirring. Subsequently the samples were centrifuged at 2500 rpm, brake rate 0 for 20 minutes. Afterwards the sodium polytungstate with the floating material was decanted into a (target) test-tube. This procedure was repeated once. 7. The sodium polytungstate was recovered by diluting it with distilled water. The test- tubes were topped up with distilled water and centrifuged at 2500 rpm brake rate 9 for 15 minutes. The diluted sodium polytungstate was recovered by decanting it in glass measuring jug. For the target test-tubes (containing the float) this procedure was repeated 3 times, for the residue 2 times. 8. The float was pipetted from the target test tube on microscope slides where the water was condensated. Subsequently it was covered with a drop of glycerin and a cover glass. Whereupon the samples were ready for optical identification. 9. Tephra shards were identified on the basis of their behavior under cross-polarized light (tephra/glass shards remain dark under cross-polarized light). Furthermore, the colour and shape of the shards were taken into account.

Lateglacial and early Holocene tephra possibly deposited in Ireland There is still no tephra-chronological framework for Ireland. Although tephras have been found in northern Ireland, their origin could not be confirmed yet (Lowe et al., 2004). Scientists are still discovering microtephra at more distant locations (Hall and Pilcher, 2002), which favors the opportunity of finding these tephras in Ireland. However, an increasing number of tephras with similar geochemistry and in close stratigraphical and temporal proximity are being detected, 45 which makes the site-to-site correlation of tephras more complex (Pyne-O’Donnell, 2007). The most likely tephras to be deposited at the study site (those which are found in the region near Ireland) are presented here (see figure 23). As time was a limiting factor, only the lower 55 cm (samples 1 to 73) and the assumed Younger Dryas (GS-1) (samples 107 to 116) were investigated for the presence of tephra; this limited the possible tephras which could be found to the Dimna, Borrobol, Penifiler and the Vedde tephra.

ka b. 2000 AD 11,0 HO Ashik (ref. 3)

12,0 Vedde (ref. 4) YD

13,0 Laachersee (ref. 4)

B Penifiler (ref. 3) 14,0

B Borrobol (ref. 2)

15,0 Dimna (ref. 1) -44 -42 -40 -38 -36 -34 NGRIP delta18O (per mille)

Fig. 23: δ18O curve for the Lateglacial derived from the NGRIP ice core with expected microtephra layers in Ireland (δ18O curve based on data from NGRIP members, 2004, Rasmussen et al., 2005, Rasmussen et al., 2006. Microtephra data based on: ref.1: Koren et al., 2008, ref.2: Pyne-O’Donnell et al., 2008, ref.3: Pyne-O’Donnell, 2007, ref.4: Turney et al., 2006).

5. Results

5.1 Climate reconstruction

5.1.1 Lithostratigraphy In the transect (figure 24) the position of the sediment core (LU-A) is indicated. From the transect an overall sediment sequence is observed of (from top to bottom): peat, gyttja, calcareous gyttja cross-cut by clay(ey) layers and ending at the glacial till with a clay layer. From the transect it is clear that the core was not taken at the place with the thickest lake infill, the reason for this is that the Lateglacial sequence (the alternation of calcareous gyttja with clay(ey) layers) was more complete at the location where the sediment core is taken. The reason for the absence of the clayey calcareous gyttja layer in the deepest part of the lake infill is thought to be the limited amount of clastic inwash of the clayey sediment not reaching the deepest part of the lake. The overall lithostratigraphy of the sediment core is described in table 3.

Depth Type of below GW Colour Remarks Period associated with deposit (cm) Vague laminations Light grey- Calcareous Sediment contains Early 610.7 - 686 brown gyttja shell and plant Holocene

remains Upper limit is gradual Younger 686 – 691.6 Silty clay Grey in two centimeters. GS-1 Dryas Lower limit is sharp.

Vague laminations 691.6 – Calcareous Light Sediment contains Allerød GI-1a t/m 1c 747.5 gyttja brown-grey shell remains.

Clayey Upper and lower limit Grayish Older 747.5 - 749 calcareous gradual in one GI-1d green Dryas gyttja centimeter. Sediment becomes more grey-green in Calcareous Light 749 - 770 colour downwards. Bølling GI-1e gyttja brown-grey Some black spots in sediment Laminations 770 - 786 Silty clay Grey Upper limit gradual in Pleniglacial GS-2 two centimeters

Table 3: Lithostratigraphy of the LU-A sediment core

From the lithostratigraphy two main things are evident. There is an apparent alternation of lake marl deposits and clayey deposits ánd the clayey deposits are relatively thin compared to the lake marl deposits. Lake marl and clay have different origins. Lake marl is formed by the utilization of carbon dioxide for photosynthesis and the resulting supersaturation of calcium carbonate in the water column; it typically forms during summer photosynthesis when biological activity is highest (Leng and Marshall, 2004). Summer temperature has a strong controlling effect on the formation of lake marl, therefore warm periods (interstadials) are characterized by the formation of lake marl. During cold periods (stadials) the summer temperature is much lower and erosion rates are much higher, because there is less vegetation cover and more mechanical weathering. Therefore the deposition of clastic materials gets more important during stadials. The thickness differences of the lake marl and the clay(ey) deposits are probably caused by higher accumulation rates of precipitation of lake marl during warm periods; during colder periods precipitation of lake marl can even cease completely (accumulation rates are further discussed in chapter 5.1.3).

stigated sediment is core indicated with LU-A. Fig. 26: Coring western transectThe Ireland. Lurga, inve

5.1.2 Loss on Ignition To estimate the biological productivity during the Lateglacial the organic matter content is determined using the loss of ignition method. In figure 25.b the results of the loss of ignition method are presented. The LOI-values vary between 7 and 28 percent of organic matter content. There is an obvious fluctuating trend visible in the graph which matches the different samples series. This fluctuating trend could be caused by differential burning of organic matter depending on the position in the furnace. Secondly, carbonates present in the sample could be partially combusted during the LOI method which could result in an additional loss on ignition; which results in an unreliable estimated percentage of organic matter content. To test how partially burning of carbonate influence the LOI-results, an experiment was set up. First a crucible with pure calcite powder was placed in the furnace to test if any of the calcite was combusted at 550 ºC for four hours. As this was the case, the next experiment was set up with a sample series of calcite powder to test how the position in the furnace influenced the percentage of combusted calcite. The percentage of calcite which was burned varied from 0 percent in the front of the furnace to 35 percent in the back of the furnace (figure 26). Clearly, the position in the furnace influenced the LOI-results dramatically. Using the spatial distribution of the relative amount of calcite burned, the LOI-results could be corrected (figure 27.c). The following correction is applied:

LOIcorrected = LOI – (CaCO3 sample * (CaCO3 position furnace/100)) [%]

LOI = LOI550 (for formula see methods LOI chapter 4.1.2)

CaCO3 sample = the volumetrically measured percentage CaCO3 in the sample

CaCO3 position furnace = the percentage CaCO3 combusted at the position in the furnace of the sample (see figure 26)

There are several peaks in the corrected organic matter curve, of which two correspond with the top two clay(ey) layers. However, for a large part of the curve the organic matter content is (below) zero, which is very unlikely because there is always some organic matter present in sediment. It seems that the amount of calcite being partially combusted is overestimated as many values are below zero. A possible explanation for this can be that in the powder form, relatively more calcite is burned than in the cohesive mixed sediment sample. However, after uniformly lowering the relative amount of calcite being partially combusted, the shape of the LOI-curve did still not match the LOI-curves from previous research and therefore this result is not used for further analysis. An alternative method to remove the organic matter content is to apply hydrogen peroxide. 21 Samples were taken on which the organic matter was removed with hydrogen 50 peroxide (see figure 25.d). The percentage of organic matter content determined with the hydrogen peroxide method range from 0 to 18 percent, with a peak organic matter content at the clayey layer around 745 cm depth.

a b c d 626 Calcareous gyttja

646 Clay

Clayey calcareous gyttja 666

686

706 Depth (cm)

726

746

766

786 0102030 0 5 10 15 20 0 5 10 15 20 Organic matter with Lithology LOI (%) LOI corrected (%) peroxide (%) Fig. 25: a) lithology of the studied sediment core. b) the organic matter content in percentages derived from the loi- method. The red lines define the different sample series which were placed in the oven for combustion. c) the organic matter content corrected for carbonates partially burned in the oven dependent on their position. After correction many values were below zero, as this is not possible these were set to zero percent. d) Organic matter content determined with hydrogen peroxide. Only 21 samples were treated with hydrogen peroxide.

18 35 15 27 18

25 23 25 22 31

33 27 23 28 32

21 22 18 28 26

2 9 5 8 0

Fig. 26: Percentage of calcite burned after combustion at 550 ºC for four hours. The top of the figure is the back of the furnace and the bottom part of the figure corresponds with the front of the furnace.

5.1.3 Isotope analysis

Carbonate content The results of the Scheibler method to estimate the percentage calciumcarbonate present in the samples is showed in figure 27. The effect of atmospheric variation measured was negligible and therefore not corrected for. The calciumcarbonate content was transferred from measured gas distance to calciumcarbonate per volume sample, by using a linear relation between known volumes of calciumcarbonate and measured gas distance.

626

646

666

686

706 Depth (cm) Depth

726

746

766

786 050100

Lithology CaCO3 (%)

Fig. 27: Measured calciumcarbonate percentages (LU-A)

The overall calciumcabonate percentage is relatively high throughout the sampled sediment core, reaching values up to almost one-hundred percent. Lower calciumcarbonate percentages coincide with clay(ey) layers which are presumed to be colder periods. The lowest calciumcarbonate percentage occurs at a depth of 691 cm at the presumed Younger Dryas (GS-1).

Oxygen isotopes The results from the oxygen isotope analysis are presented in figure 28.

686

706 Depth (cm) Depth

726

746

766

786 -8,00 -6,00 -4,00 -2,00 Lithology LU-A δ18O (per mille)

Fig. 28: Oxygen isotope result for the study site Lurga

There are two distinct periods with lower δ18O values of around -7; from 786 to 770 ánd from 691 to 689 cm; which are thought to correspond with respectively the Pleniglacial (GS-2) and the Younger Dryas stadial (GS-1). From 770 to 691 cm the δ18O values are relatively high, values fluctuate around -4 δ18O per mille. Around 748 and 705 cm somewhat lower δ18O values are recognizable. The period from 770 to 691 is thought to correspond with the Bølling-Allerød Interstadial. The LU-A δ18O results seem to correspond very well with the NGRIP δ18O curve (this will be further examined in the discussion). As the different Lateglacial oscillations are clearly recognizable in the δ18O curve of the study site, estimated age and resulting accumulation rates are derived by wiggle-matching it to the NGRIP δ18O curve (figure 29 and 30). The δ18O curve of the study site is wiggle matched to the NGRIP δ18O curve by shrinking and stretching parts of the curve, while distinct peaks and valleys in the δ18O curve of the study site are assigned estimated ages derived from known NGRIP ages of matching maxima and minima. It should be stressed that the derived ages for the study site are estimated. With these estimated ages in relation to depth, accumulation rates can be assessed. During stadials (YD and Aegelsee oscillation) accumulation rates are noticeably lower than during the interstadials; which, as discussed in chapter 5.1.1, is caused by relatively higher depositional rates of precipitating marl than depositional rates of clastic sediment input. 53

Wiggle matching δ18O curves Accumulation rates

11600 11500 -2,00

Accumulation rates d18O curve

12100 12000 NGRIP -3,00

Study site

12500 12600 -4,00

13000 13100 -5,00 18O δ 13500 Age (yrs b2k)

13600 Estimated age (yrsb2k) -6,00

14000

14100 -7,00 14500

14600

15000 -8,00 670,00 690,00 710,00 730,00 750,00 770,00 790,00 Depth (cm) Fig. 29: Wigglematched δ18O curve of the study site (from Fig. 30: Estimated accumulation rate at the study site a depth of 770 cm) to the NGRIP δ18O curve compared to the δ18O curve of the study site

Carbon isotopes Carbon isotope ratios are also determined using mass spectrometry. The results are presented in figure 31.

686

706 Depth (cm)

726

746

766

786 1,00 3,00 -8,00 -6,00 -4,00 -2,00 13 Lithology δ C (per mille) δ18O (per mille)

Fig. 31: LU-A δ13C results of the study site compared to the LU-A δ18O results of the study site

The LU-A δ13C values range between 1,20 and 3,90 δ13C per mille. From 786 to 771 cm the δ13C values are quite constant at 2,80, after which they begin to rise to 2,90 at 753 cm. From 753 to 711 cm δ13C values gradually decrease to 1,20. From 711 to 700 cm a small increase in δ13C values is visible to 1,95, after which they decrease again to 1,25 at 692 cm. At this depth δ13C values suddenly increase to 3,10 at 690 cm after which δ13C values decrease again. Between 765 to 695 the LU-A δ13C values fluctuate around a mean trend in contrast to the bottom and top 15 cm of the sampled core. When the LU-A δ13C result is compared to the LU-A δ18O curve, the result can be best described in three different parts (see grey and white blocks in figure 37). First, in the lower part of the sampled core (from 786 to 745 cm depth) both curves have approximately the same shape. From 745 to 715 cm δ13C values gradually decrease while δ18O values stay approximately constant. From 715 cm depth both curves show a high negative correlation at the supposed Younger Dryas stadial (GS-1) from 695 to 680 cm. As expected, measured δ13C values are very high as a consequence of dissolution of carbonates from the limestone bedrock. The main controlling factor on fluctuations of the δ13C in the lake water is normally assumed to be the biological productivity which depends on summer temperatures. However, this seems not always the case at the study site; in the discussion this will be further elaborated.

5.2 Chronological Framework

5.2.1 Tephra In figure 32 the number of, optically identified, potential tephra shards is presented.

686

706 Depth (cm)Depth

726

746

766

786 0102030 Lithology Potential tephra shards

Fig. 32: Graph with potential tephra shards found in core LU-A

At a depth of 750 cm, which coincides with the onset of the Aegelsee oscillation (GI-1d), a shard population has characteristics of tephra shards, with a peak shard count of eleven. The shards are relatively thick and platty, have a grayish transparent colour and sharp edges (see figure 33). However, this population is not thought to be tephra as the shards are relatively thick and irregular, furthermore experts in the tephrochronology have determined (on the basis of photographs) that the shards were not tephra (through personal communication of N. van Asch). At a depth of 743 (and 739) cm a peak in potential tephra shards is identified (25 (and 7)). The shards are relatively thin and platty and have vague edges. These shards are thought not to be tephra (through personal communication with dr. W.Z. Hoek). In the samples were sometimes minerals present which had a density of more than 2.5 g/cm3. With the density separation technique all sediment which has a density of more than 2.5 g/cm3 should have been removed. However, in these samples were also biogenic silica present; it is thought that the biogenic silica have attached to the minerals, which lowered their average

56 density to less than 2.5 g/cm3. Furthermore, during the optical identification of tephra shards, some thick, grayish shards were discovered, which showed hardly any signs of erosion (figure 34).

Fig. 33: Shard population found at the study site in core LU-A at a depth of 750 cm.

Fig. 34: Glass shard found during optical identification of tephra shards, thought to originate from (with glass shards) contaminated cover glass.

These shards were present in the sample ánd on top of the cover glass. The cover glasses used were examined and contained shards on it probably formed during fabrication. These cover glasses were not used for further analysis and the samples already prepared with these cover glasses were re-sampled. During this study no tephra was found. The reason for this can either be, that there was something wrong with the method used or that there was no tephra present in the samples (this will be further explored in the discussion). 57

6. Discussion

6.1 Organic matter content

Both the position in the furnace and the partial combustion of carbonates, influenced the results of the organic matter content in such a way that the LU-A LOI result was assumed to be unreliable. A possible solution to the partial combustion of carbonates is, to lower the combustion temperature to a temperature where no more carbonates are combusted. The disadvantage of this method is, that it is possible that part of the organic matter is not ashed completely. Furthermore during combustion, the row samples closest to the door of the furnace significantly deviated below the average percentage of calciumcarbontes burned (figure 26); which is thought to be caused by the door of the furnace not closing completely. The effect of the bad closure can be decreased by not using the front row of the furnace. However, the best solution would be to fix the door or purchase a new furnace. As the combustion of organic matter applied in this study did not yield a reliable (corrected) LOI curve; the organic matter content, optained with the use of hydrogen peroxide, is further examined. LOI-curves of previous research in the area of the study site are all quite similar in shape, in figure 35 the LOI-curve of an earlier research at the study site Lurga is presented (Paus et al., 1994) and compared to the organic matter content result from this study. Both organic matter contents (from this study and a previous study at Lurga) have roughly the same average percentage (approximately five to ten percent). However the LU-A organic matter content has a peak percentage of 18 percent at approximately 750 cm depth, while the other curve does not show any obvious increase in organic matter content at this depth. The LOI-curve from the previous research at Lurga (Paus et al., 1994) has two peak organic matter percentages at the bottem and top clay layer in its lithology; these high organic matter contents are not recognizable in the LU-A result. The reason for the lack of peak organic matter perctages at the study site, can be that only 21 samples of the LU-A core are examined for organic matter content; and therefore these peak percentages could have been missed. Further research is needed into the use of hydrogren peroxide to determine organic matter content, to establish more qualitatively its value for climate reconstruction.

Fig. 35: Organic matter content from the study site (LU-A) compared to LOI results from a previous research at Lurga (Paus et al., 1994), both records are correlated with their lithology.

6.2 Isotope analysis

Carbonate content When the LU-A calciumcarbonate result is compared to the NGRIP-δ18O curve, the shape of the curves are quite similar; indicating that the carbonates in this study are a good qualitative measure for summer temperatures and that the carbonates are most likely to be almost entirely authigenic in origin. When compared to previous research in the region (figure 36), the shape of the carbonate curve at Lurga matches the carbonates curve from Tory Hill quite well. The percentages of carbonate in the sediment cores are in the same order of magnitude and the overall shape of the curve is the same, although the presumed Younger Dryas stadial (GS-1) is much more prolonged at Tory Hill than at Lurga. 59

yrs b2k 11,0

12,0

13,0

14,0

15,0

0 50 100 -44 -39 -34

Tory Hill CaCO (%) LU-A CaCO (%) NGRIP δ18O (per mille) 3 3 Fig. 36: Carbonate percentages (of this study) compared to the NGRIP δ18O and to a carbonate curve from a study at Tory Hill (O'Connell et al., 1999). Matching maxima and minima are connected with (red) dotted lines.

Oxygen isotopes In figure 37 LU-A δ18O curve is compared to the NGRIP δ18O curve; the shape of the curves resemble each other very well. Besides the large-scale (stadial-interstadial) oscillations, even small oscillations (e.g. Gerzensee (GI-1b) and Aegelsee oscillation (GI-1d)) are visible in the LU-A δ18O curve. As the resemblance between both graphs is that high, the same forcing on their δ18O values is expected. As explained in the method chapter changes in the temperature at the source of the precipitation (the ocean) can explain the high correlation between both curves. It is therefore assumed that the LU-A δ18O curve at the study site (and at the NGRIP study site) is mostly forced by the δ18O of precipitation, however also local temperature is expected to have influenced the LU-A δ18O of the lake water by reinforcing or weakening the δ18O signal of the precipitation .

yrs b2k 11,0 666

686 12,0 YD

706 Ge rzens Depth (cm) ee o scillat ion 13,0 726

746 Aegelsee oscillation 14,0

766

15,0 786 -8,00 -6,00 -4,00 -2,00 -44 -39 -34 Lithology δ18O (per mille) NGRIP δ18O (per mille)

Fig. 37: Oxygen isotope result compared to the NGRIP δ18O curve. Matching peaks and valleys are connected with dotted (red) lines.

As there is a high resemblance between the LU-A and NGRIP δ18O curve, oscillations visible in the LU-A δ18O curve are assumed to correlate with the oscillations known from the NGRIP δ18O curve. The main difference between the two curves is, that the δ18O curve from the study site does not show a downward trend in the δ18O values towards the Younger Dryas. These persisting high δ18O values during the Bølling-Allerød Interstadial are also visible in another δ18O records from the area (figure 38). However, the δ18O curve of Red Bog (which is situated very close to Lough Gur) does show a downward trend towards the Younger Dryas similar to the NGRIP δ18O curve. As both areas are situated very close to each other, local temperature cannot explain the differences between both graph. Therefore the LU-A δ18O results cannot be directly interpreted as temperature, neither local temperature or temperature at the ocean. However, small scale oscillations can be clearly identified and correlated to other Lateglacial terrestrial, marine and ice core records. Therefore it is presumed that the persisting high δ18O values at the study site are caused by local phenomena (f.i. partial closure of lakes) which determines the average trend in δ18O values but does not influence smaller oscillations.

a) b)

d) yrs b2k 11,0

12,0

13,0

14,0

15,0

-8,00 -6,00 -4,00 -2,00 -44 -39 -34

18 δ O (per mille) δ18O (per mille) δ18O (per mille) δ18O (per mille) Red Bog Lough Gur Lurga (this study) NGRIP Fig. 38: Oxygen isotope results for a) Red Bog b) Lough Gur (Ahlberg et al., 1996) c) Lurga (this study) and d) the NGRIP ice core. Matching peaks and valleys are connected with dotted lines.

Lurga δ18O (per mille)

-2,00

-4,00

-6,00

-8,00 Summer SST (ºC) SST Summer

Fig. 39: Sea surface and salinity estimates for core NA 87-22 off the west-coast of Ireland (Duplessey et al., 1996). Results are compared to the δ18O result of this study. 62

The δ18O record of the study site is also compared to the estimated summer-SST (summer sea surface temperatures) of the marine core offshore of Ireland (NA 87-22) (figure 39). Although the marine record is not as detailed, both graphs show resemblance and don’t show a downward trend towards the Younger Dryas as the NGRIP δ18O values did. This implies that the Lateglacial climate at the site of the marine core and the study site was approximately the same.

Carbon isotopes The variations in the LU-A δ13C curve (in correlation to the LU-A δ18O curve) (see figure 31) are discussed and interpreted in more detail below. From 786 to approximately 745 cm depth the δ18O and δ13C curves have approximately the same shape. This similarity in shape can be explained if both the δ18O and δ13C values are forced by local summer temperature. This means that the main controlling factor of the δ13C values in this part of the curve was the biological productivity. The decreasing downward trend in δ13C values, from approximately 745 to 715 cm depth, is difficult to explain. The decreasing trend in LU-A δ13C values corresponds with prolonged high LU-A δ18O values in comparison the the NGRIP δ18O curve, which is thought to be caused by a local phenomina. A possible explanation would be that δ13C values decrease when the area changes from exposed limestone bedrock to a vegetated surface. However, during an earlier research at Lurga pollen analysis was done for the Lateglacial and this showes that the exact opposite is happening (figure 40), the vegetation is changing from relatively closed vegetation to a more herbaceous vegetation (based on correlation through lithology the period 745 to 715 cm in LU-A corresponds with 250 to 200 cm depth in figure 40). Therefore, afforestation cannot be the reason for the decrease in LU-A δ13C values.Furthermore, iIt is not likely that these decreasing δ13C values are caused by a decrease of the input of allogenic carbonate, because if (a part of) the carbonate would be allogenic in origin, it would be accompanied by a clastic input which is (almost) absent at this part of the curve. An appropriate explanation for the decline in LU-A δ13C values is not found and should be further explored. From approximately 710 to 680 cm depth the δ18O and δ13C curves show a high negative correlation. The reason for this high correlation cannot be that the lake was hydrologically closed because if that would be the case there would be much more variation in the δ18O and δ13C values (as they would be mostly influenced by changes in precipitation) and the δ18O curve would not match the NGRIP δ18O curve that well. An explanation for this synchronism could be, that during this period there was an input of allogenic carbonate, which changed the δ18O and δ13C values to isotope values which are similar to detritus (mainly limestone in origin). Furthermore, Diefendort et al. (2008) propose that changes in the aquatic productivity, aquatic respiration and algal isotope fractionation can explain these shifts in δ13C.

Fig. 40: Composite pollen curves from a previous research at the study site Lurga (Paus et al., 1994). The lithological silty gyttja units (matching approximately the local PAZs 2 and 5) are thought to correspond with the top two clay(ey) layers in LU-A.

The carbon isotopes ratios at Lurga are compared to previous research at Tory Hill (figure 41).The results are difficult to compare as the variation in the δ13C values is very small. However, the results are of the same magnitude and mainly in the lower part of the curves the shape of the curves is quite similar. In the higher part of the curves the δ13C values at Lurga decrease, while the δ13C values at Tory Hill fluctuate. Concluding, δ13C values are of the same magnitude and there might be some resemblance in the shape of the curve, however the variation in δ13C values is to small to properly compare both graphs.

1234

δ13C Lurga δ13C Tory Hill (per mille) (per mille)

Fig. 41: δ13C values of the study site compared to a δ13C record from a previous study at Tory Hill (O’Connell et al., 1996)

6.3 Tephra

In this study no tephra was found. The reason for this can either be, there was no tephra present in the samples or that there were errors made in the method of tephra extraction. The absence of deposited tephra can be caused by three reasons. Firstly, the spread of the ash in the atmosphere could have been such that it has not crossed the study site. Secondly, it is possible that tephra is deposited at the study site but did not accumulate at the exact coring location. Finally, tephra could have been deposited but is not present any more because of natural glass corrosion in the sediment. As the environment at the study site during the Lateglacial was highly alkaline, the most likely type of glass corrosion to have taken place is network dissolution (of sioxane bonds); which can lead to complete dissolution of tephra shards (Blockley et al., 2005). It is questionable if the absence of tephra at the study site is the reason for

65 not finding tephra during this study. As Lateglacial tephras are found in Northern-Ireland (Lowe et al., 2004), Scotland (Turney et al., 1997, Pyne-O’Donnell, 2007) and throughout a large part of Europe, it is very likely that tephra has been deposited at the study site. There are procedures in the method used, which could have caused the absence of tephra. To remove carbonates, a 5% HCL solution was added and left to react overnight. From research it is shown that treatment with acidic substances can decrease the number of shards (Blockley et al., 2005). Furthermore, to remove any clay in the samples, the samples were sieved over a 31 μm sieve. Any tephra smaller than 31 μm was therefore lost. And finally, to recover the sample on a cover glass, the descended sample was pipetted from the target test tubes on mounts where the water was evaporated. When the water with the sample was pipetted, the sample was left to descend in the pipette for approximately 3 minutes. Afterwards two droplets of the water with the sample were pipetted on the cover glass. The time for descending of the sample in the pipette could have been too short, therefore the sample with possible tephra could not have been descended far enough to be included in the two droplets which were pipetted on the cover glass. It would be better to pipet the whole sample from the target test tubes on to a cover glass, instead of two droplets; this way no tephra can be missed because the whole sample was not examined. It would improve the efficiency and reliability of future tephra research if the researcher would, a priori, be familiarized with different types of tephra. In that way, other types of glass shards (than tephra) would not be regarded as being tephra as easily.

6.4 Climate reconstruction

In figure 42 the results, of the climate proxies used, are presented (except for the organic matter content). The δ13C values only represent temperature change in the lower forty centimeters of the core. As the studied sediment core is thought to cover the Lateglacial period, it is correlated to the NGRIP event stratigraphy through wigglematching. However, 14C-dating should be performed to construct a more reliable chronology. All climate proxies show a uniform climate signal during the Lateglacial in which the following periods can be distinguished: Pleniglacial, Bølling, Aegelsee oscillation, Allerød, Gerzensee oscillation, Younger Dryas and the onset of the Holocene.

626 ne ooeeYD Holocene Onset 646

666

686

706 Gerzensee leø Bølling Allerød Depth (cm) Depth

726

746 AO

766 Oldest Dryas

786 0 50 100 -8,00 -6,00 -4,00 -2,00 1,00 3,00 Lithology Carbonates (%) δ18O (per mille) δ13C (per mille)

Fig. 42: Multiproxy analysis on sediment core LU-A at the study site Lurga. Results include lithology, carbonates, oxygen and carbon isotope ratios.

Parallel to this study, Arthur Lutz has performed a chironomid analysis on the LU-A sediment core. In figure 43 the results from the chironomid analysis are compared to the LU-A δ18O results from this study. The spread in the chironomids DCA (detrended correspondence analysis) graph shows how much the chironomid assemblages differ from each other. Since chironomids are a proxy for summer temperature, it is assumed that the spread in the graph indicates how much the summer temperatures vary between the samples, without actually assigning absolute temperature values to these variations. From figure 43 it is clear that the LU-A δ18O result and LU-A chironomid result correlate very well. There are some minor leads and lags visible, but large- and small-scale oscillations are clearly visible in the chironomid result and are easily correlated to the LU-A δ18O result. To even better correlate the chirnomid results to other results in this study, more samples should be analysed.

666

686

706 Depth (cm)

726

746

766

786 -8.00 -6.00 -4.00 -2.00 3 2 1 0 18 Choronomomids DCA (SD units) Lithology δ O (per mille) Fig. 43: LU-A δ18O results compared to results for a chronomid detrended correspondence analysis (DCA in standard deviation units) on the LU-A sediment core (A.F.Lutz, 2009)

7. Conclusions

The Lateglacial was a period of relative rapid climate fluctuations, which is an interesting period to study as it gives insight in the variability of the climate. To reconstruct Lateglacial climate and timing at the study site, several climate proxies were examined and microtephra analysis was performed to construct a chronological framework.

The research questions formulated are answered below:

1. How did climate change during the Lateglacial at the study site?

At the bottom of the sediment core the Pleniglacial (GS-2) is preserved (see figure 42), which is characterized by clastic sediments and low temperatures. After the Pleniglacial, climate changed to a warm period (the Bølling Interstadial (GI-1e)) during which calcium carbonates are precipitated in the lake. This warm climate was briefly interrupted by the Aegelsee oscillation (GI- 1d) in which the precipitated carbonates are accompanied by a more clastic input of clay. After this brief cooling δ18O values rise again and the sediment is again made up almost entirely by calcium carbonates. This period is called the Allerød Interstadial (GI-1a-c), during the Allerød two cold phases are recognizable from the δ18O curve; of which the youngest one is called the Gerzensee Oscillation (GI-1b). After this prolonged warm period δ18O values drop abruptly to start a cold phase (the Younger Dryas Stadial (GS-1)) during which the sediment is almost completely made up out of clay, the Younger Dryas is followed by the onset of the Holocene.

2. Is tephra deposited at the study site and if so, can shards be appointed to a known tephra?

From tephra analysis on the sediment core it is apparent that no tephra was present. The absence of tephra can either be explained by non-deposition of tephra at the coring site, or by an error in the tephra extraction method.

3. How do the results of the multiproxy analysis correlate to previous research in the region?

The results of the multiproxy analysis correlate very well to previous research in the region. The shape and magnitude of the carbonate concentrations is very similar to a more southerly record at Red Bog (figure 36). The oxygen isotope curve correlates well to other curves from the region (see figure 38), however δ18O values stay high during the Bølling-Allerød Interstadial which can be caused by a local phenomena. The δ18O correlates well to a nearby marine sediment record (figure 39) and the NGRIP ice core record (figure 37). The carbon isotope ratios from Lurga are in

69 the same magnitude as results from a record at Tory Hill (figure 41), the variations in the δ13C values are too small to comment on the similarity (or dissimilarity) in shape of both curves.

4. How do the results from this study correlate to results from ice and marine cores?

The δ18O curve at the study site matches the NGRIP ice core record very well. The main difference is that the δ18O values of the NGRIP record show a downward trend during the Bølling- Allerød Interstadial, in contrast to the δ18O values at the study site which stay fairly constant throughout this period. When compared to a marine core, taken off shore of western Ireland, the δ18O curve of the study site matches the reconstructed summer-SST of the marine core very well; indicating that the climate record at the study area provides a good analogue for the maritime environment of the North-Atlantic ocean.

Acknowledgements

During this study Nelleke van Asch and dr Wim Hoek have acted as my supervisors. I would like to thank them for their assistance with the fieldwork in Ireland and for their supervision and encouragement throughout this study. Furthermore, I would like to thank Arthur Lutz who did a seperate project on this study; but has always taken the time to listen and give advice on my project.

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