MASARYK UNIVERSITY
FACULTY OF SCIENCE
DEPARTMENT OF GEOGRAPHY
SROVNÁNÍ VÝVOJE SEVERU A JIHU – HOLOCENNÍ ABIOTICKÁ
PALEOLIMNOLOGIE VYSOKÉ ARKTIDY A PŘÍMOŘSKÉ
ANTARKTIDY
COMPARING THE EVOLUTION OF NORTH AND SOUTH – HOLOCENE
ABIOTIC PALAEOLIMNOLOGY OF HIGH ARCTIC AND MARITIME
ANTARCTICA
Master’s diploma thesis
Tomáš Čejka
SUPERVISOR: Assoc Prof. Daniel Nývlt, PhD. BRNO 2018
Bibliografický záznam
Autor: Bc. Tomáš Čejka
Přírodovědecká fakulta, Masarykova univerzita
Geografický ústav
Název práce: Srovnání vývoje severu a jihu – holocenní abiotická
paleolimnologie Vysoké Arktidy a Přímořské Antarktidy
Studijní program: Geografie a kartografie
Studijní obor: Fyzická geografie
Vedoucí práce: doc. Mgr. Daniel Nývlt, Ph.D.
Akademický rok: 2017/2018
Počet stran: 88+6
Klíčová slova: jezerní sedimenty, Neoglaciace, klimatické optimum, Malá
doba ledová, ostrov Vega, Západní Špicberk,
paleolimnologie, magnetická susceptibilita, rentgenová
fluorescence, uhlík, síra, proxy záznam, rozsivky, 14
C, 210
Pb,
137
Cs, časově-hloubkový model
Bibliographic Entry
Author: Tomáš Čejka, BSc
Faculty of Science, Masaryk University
Department of Geography
Title of Thesis: Comparing the evolution of North and South – Holocene
abiotic palaeolimnology of High Arctic and Maritime
Antarctica
Degree Programme: Geography and Cartography
Field of Study: Physical geography
Supervisor: Assoc Prof. Daniel Nývlt, PhD.
Academic Year: 2017/2018
Number of Pages: 88+6
Keywords: lake sediments, Neoglaciation, climate optimum, Little Ice
Age, Vega Island, Spitsbergen, palaeolimnology, magnetic
susceptibility, X-ray fluorescence, carbon, sulphur, proxy
record, diatoms, 14
C, 210
Pb, 137
Cs, age-depth model
Abstract
The lake systems in the Polar regions have experienced pronounced environmental changes
during the Late Holocene, and most of these were appropriately reflected in the lake
sedimentary records. For this thesis, two cores from the Vega Island (Lake Anonima,
Antarctic Peninsula) and the Spitsbergen (Lake Mathiesondalen 3, Svalbard) were analysed
using multi-proxy approach to obtain more-detailed view on the palaeoenvironmental
evolution. Geochemical (carbon and sulphur content, X-ray fluorescence spectroscopy),
petrophysical (magnetic susceptibility, grain-size composition) and biological (diatom
biostratigraphy) analyses were performed, and supplemented by a robust sedimentation
chronology based on radiocarbon (14
C) and short-lived (210
Pb, 137
Cs) isotope dating.
Principal component analysis (PCA) and cluster analyses were applied to access the
dominating patterns in the sediment composition. Lake Anonima covers the last 2400 cal.
yrs b2k and provides a record of the climate history in the James Ross Archipelago, Northeast
Antarctic Peninsula. The Lake Anonima records the Late Holocene Hypsithermal and
the Neoglacial Period. The Lake Mathiesondalen 3 covers the last 700 cal. yrs b2k and
presents a record of changes in the Mathiesondalen lake and catchment, as well as unique
perception of the Little Ice Age mild course.
Abstrakt
Jezerní systémy v polárních oblastech prodělaly během pozdního holocénu řadu výrazných
environmentálních změn, které lze vhodně rekonstruovat z jezerních sedimentárních
záznamů. Pro tuto práci byly analyzovány dvě jádra z ostrova Vega (jezero Anonima,
Antarktický poloostrov) a Západního Špicberku (jezero Mathiesondalen 3, Svalbard) a to
s využitím multi-proxy přístupu za účelem získání detailnější informace o
paleoenvironmentálním vývoji. Byly aplikovány geochemické (podíl uhlíku a síry,
rentgenová fluorescence), petrofyzikální (magnetická susceptibilita, laserová granulometrie)
a biologické (biostratigrafie rozsivek) analýzy, jež byly posléze doplněny o robustní
chronologii sedimentace založenou na radiouhlíkovém (14
C) datování a datování izotopů
s krátkým poločasem rozpadu (210
Pb, 137
Cs). Za účelem zjištění dominantních trendů
v sedimentárním složení byla provedena, společně se shlukovou analýzou, analýza hlavních
komponent. Sedimentární záznam z jezera Anonima zahrnuje posledních 2400 kal. let b2k
a poskytuje záznam klimatické historie souostroví Jamese Rosse. Jezero Anonima
zaznamenává pozdně holocenní klimatické optimum a období neoglaciace. Záznam z jezera
Mathiesondalen 3 zahrnuje posledních 700 cal. let b2k a zahrnuje změny v jezeře a jeho
povodí, stejně tak jako unikátní záznam mírnějšího průběhu Malé doby ledové.
Acknowledgements
Here, I want to express my sincere gratitude to thesis supervisor Daniel Nývlt for his
inspiring, encouraging and professional attitude. His valuable commentary contributed to a
significant improvement in the whole work, as well as did his supporting spirit. I must
acknowledge a foundation of my scientific interest has been, definitely, set by him.
The work on this thesis has been supported by the projects of the Ministry of Education,
Youth and Sports of the Czech Republic no. LM2015078 and
CZ.02.1.01/0.0/0.0/16_013/0001708.
I would also like to show gratitude to many other colleagues, namely Lenka Lisá (Institute
of Geology, Czech Academy of Sciences) for providing the laser granulometry analysis and
for her cooperation during the measurement, Kateřina Kopalová and Marie Bulínová
(Department of Ecology, Charles University) together with Bart Van de Vijver (Antwerp
University) for diatom biostratigraphy analysis and discussions, Viktor Goliáš (Institute of
Geochemistry, Mineralogy and Mineral Resources, Charles University) for short-lived
radioisotope dating, Eva Geršlová (Department of Geological Sciences, Masaryk University)
for providing X-ray fluorescence method and Matěj Roman (Department of Geography,
Masaryk University) for his assistance and help. I am also grateful to Jan Kavan (Department
of Geography, Masaryk University), Silvia H. Coria, Marcos A. E. Chaparro and Juan M.
Lirio (Instituto Antártico Argentino) for core retrieval during the field campaigns.
I also gratefully acknowledge the support from my family, and most importantly, Míša.
Thank you for showing me the right direction.
Declaration
I hereby certify that this master’s diploma thesis is entirely the result of my own work
and that I have conducted this paper under supervising of Assoc Prof. Daniel Nývlt, PhD. I
have faithfully and properly cited all sources used in this thesis.
Brno, 3rd
May 2018 ...….………………………………
Tomáš Čejka
9
TABLE OF CONTENT
1 INTRODUCTION, RATIONALE AND AIMS...............................................................11
2 LITERATURE REVIEW – PALAEOENVIRONMENTAL RECORDS .......................12
2.1 North-eastern Antarctic Peninsula .............................................................................12
2.1.1 Physical setting and geographical description.....................................................12
2.1.2 Palaeoenvironmental records...............................................................................15
2.2 Svalbard Archipelago.................................................................................................16
2.2.1 Physical setting and geographical description.....................................................16
2.2.2 Palaeoenvironmental records...............................................................................18
3 SETTINGS........................................................................................................................21
3.1 Lake Anonima, Vega Island, north-eastern Antarctic Peninsula ...............................21
3.2 Lake Mathiesondalen 3, Mathiesondalen, Central Svalbard......................................23
4 MATERIALS AND METHODS......................................................................................26
4.1 Cores retrieval, sediment sampling and description ..................................................26
4.2 Radiocarbon dating ....................................................................................................26
4.3 210
Pb and 137
Cs (short-lived radionuclide) dating.......................................................28
4.4 Signal of magnetic susceptibility ...............................................................................28
4.5 Grain-size distribution................................................................................................29
4.6 Geochemical composition..........................................................................................30
4.6.1 Lithophile elements determination by means of X-ray fluorescence (XRF) ......30
4.6.2 Carbon and sulphur content.................................................................................31
4.7 Diatom biostratigraphy...............................................................................................31
4.8 Statistical data analyses..............................................................................................32
5 RESULTS .........................................................................................................................33
5.1 Lake Anonima............................................................................................................33
5.1.1 Lithology and age-depth model...........................................................................33
5.1.2 Physical proxies...................................................................................................34
5.1.3 Geochemistry.......................................................................................................36
5.1.4 Diatom biostratigraphy (by Kateřina Kopalová and Marie Bulínová)................37
5.1.5 PCA .....................................................................................................................40
5.2 Lake Mathiesondalen 3 ..............................................................................................41
5.2.1 Lithology and age-depth model...........................................................................41
5.2.2 Geochemical and physical proxies ......................................................................43
10
5.2.3 PCA .....................................................................................................................46
6 DISCUSSION...................................................................................................................47
6.1 Lake Anonima evolution in the Late Holocene .........................................................47
6.1.1 2440–2050 cal. yrs b2k: Productive warmer period............................................47
6.1.2 2050–1520 cal. yrs b2k: Colder phase → decline in primary productivity.........50
6.1.3 1520–220 cal. yrs b2k: Filling of the lake, high diatom variability ....................51
6.1.4 220 cal. yrs b2k–recent: Recent lake development .............................................52
6.1.5 Comparison of Lake Anonima proxy data with other palaeoclimatic
reconstructions from wider area ...................................................................................52
6.2 Lake Mathiesondalen 3 evolution in the Late Holocene............................................58
6.2.1 700–380 cal. yrs b2k: High sulphur and χ and low TOC ....................................58
6.2.2 380 cal. yrs b2k–recent: Variable TOC and TC enhancement............................61
6.2.3 Comparison of Lake Mathiesondalen 3 proxy data with other palaeoclimatic
reconstructions from wider area ...................................................................................63
6.3 “Comparing the evolution of North and South since 700 cal. yrs b2k”.....................66
7 SUMMARY AND CONCLUSIONS ...............................................................................71
REFERENCES ....................................................................................................................72
APPENDIX..........................................................................................................................88
11
1 INTRODUCTION, RATIONALE AND AIMS
Polar Regions have experienced significant environmental alternations associated
predominantly with hydrometeorological and thus long-term climatic fluctuations in the Late
Holocene (4250 years b2k – modern era). These changes led to a formation of differentlyparameterized
terrestrial ecosystems, glacier zones, catchment areas, and most importantly
for this study, lake systems. In the chapter 2 of this thesis, author presents a comprehensive
literature review focused on a physical-geographical setting of selected polar areas with a
particular emphasis on the Late Holocene types of paleoenvironmental records
supplemented by actual researches. This brief summary will enable the reader to understand
later-presented (chapter 5 and 6) research results in their larger complexity and coherence.
Based on a composite and traditionally used modern multi-proxy analyses, two lacustrine
sedimentary archives from Antarctica (Lake Anonima) and Svalbard Archipelago (Lake
Mathiesondalen 3), are going to be subjected to a high-resolution, using both biotic and
abiotic, techniques, to extract information about the short- as well as long-term reactions of
these systems to driving environmental changes. The main points of this thesis aim to
provide answers to the following questions. “How have evolved the studied polar areas over
the time? Have there been recognizable shifts in particular proxy parameters? How have
been the biodiversity spectra changing? How are the glacier behaviour and hydrological
regime changes marked in the records?” And many others…
Due to the wide relationship complexity within the physical-geographical sphere, this
thesis does not aim to explain all causes and consequences in their best detail and the closest
look. In such a situation, the reader is referred to an appropriate literature discussing selected
phenomena in a wider scope.
12
2 LITERATURE REVIEW – PALAEOENVIRONMENTAL RECORDS
2.1 North-eastern Antarctic Peninsula
2.1.1 Physical setting and geographical description
The Antarctic Peninsula (AP) comprises a narrow (<250 km wide) chain of mountains, rising
to a maximum of 3500 m (but mostly much lower) along its ~1250 km length. It constitutes
an averagely continuous plateau with few disruptions below 2000 m and it projects
substantially further north than the rest of Antarctica. In many respects the AP is geologically
similar to the South American cordillera. It comprises a pre-Jurassic basement overlain and
intruded by Jurassic–Neogene magmatic arc rocks related to eastwards-directed subduction
of oceanic lithosphere along the western (Pacific) margin of the AP. A comprehensive
description of the setting of the AP is given by Domack et al. (2003b).
The AP is heavily glaciated with outlet glaciers terminating in the ocean along both its
western and eastern margins. There is a clear boundary, south of which glacier termini feed
into ice shelves, and north of which there are no ice shelves (Morris and Vaughan, 2003).
This boundary is further south (67–70°S) on the western side of the AP than on the east (64–
68°S), although during the last decades there have been recorded several ice shelves
disintegrations around the AP (Cook and Vaughan, 2010). There are only few larger ice-free
areas, where terrestrial palaeoenvironmental records from the Holocene have been preserved
(see 2.1.2), with perhaps the most important being in the north-eastern part the AP (James
Ross Archipelago), South Shetland Islands, eastern Alexander Island (Ablation Point
Massif) and the coastal area around Marguerite Bay. There are numerous small ice-free areas
on ridges between outlet glaciers, and on nunataks of the central plateau, but most of these
are rocky with few surface deposits or lakes (Bentley et al., 2009).
In many aspects, the AP is climatically different to the Antarctic continent (Fig. 1). It is
narrow piece of land thus it experiences a strong marine influence, particularly on its western
side, which is exposed to the southern westerly winds. The AP is the most northerly part of
Antarctica and as such is the most subjected to the influence from mid-latitudes. In
particular, the western side of the AP receives relatively warm, moist air masses derived
from mid-latitudes (Bentley et al., 2009). Moreover, the western AP is the only part of
Antarctica where there is a demonstrated negative correlation between winter temperatures
(at Faraday/Vernadsky Station) and the extent of sea-ice (in the Bellingshausen Sea to the
13
west of the station; King, 1994). It has been suggested that sea-ice may play an important
role in environmental change around the AP (Vaughan et al., 2003) and affects also surface
air temperatures. The eastern side of the AP is less influenced by the westerlies and
experiences a substantially colder and drier climate because of the northwards extension of
cold continental air masses from the Antarctic interior into the Weddell Sea embayment
(Reynolds, 1981). Specifically, cold barrier winds (or low-level jets; Schwerdtfeger, 1975;
Stensrud, 1996) flowing along the eastern coast of the AP bring cold and dry air towards the
northern part of the AP.
Towards marine environment, the Antarctic Circumpolar Current (ACC) flows clockwise
in a broad zone around Antarctica, carrying Circumpolar Deep Water (CDW), and is
separated from the coast by the katabatic-driven west-flowing Antarctic Coastal Current.
The north–south landmass of the AP acts as a barrier to this flow and so the southern
boundary of the ACC runs NE along the edge of the AP shelf as the ACC is deflected to the
north and through the constriction of the Drake Passage (Fig. 1). Thus, the Pacific margin
continental shelf of the AP is likely to be affected by any variations in the flow of the ACC.
This is particularly important because the ACC propagates changes in oceanic conditions
between the Pacific, Atlantic and Indian Oceans (Simmonds, 2003), including the effects of
the El Niño Southern Oscillation (ENSO; Bentley et al., 2009).
Some areas of the continental shelf on the western side of the AP experience intrusions
of CDW. This is a relatively warm (>1.5 °C), salty (34.65–34.7‰) intermediate depth water
mass (Klinck et al., 2004) that is derived from modified North Atlantic Deep Water. The top
of the CDW is at 200 m depth and so only areas of continental shelves deeper than this are
affected. These intrusions bring heat and salt onto the continental shelf. CDW is substantially
warmer than typical Antarctic surface waters and so where mixing occurs then intrusions of
CDW along the western AP shelf are characterized by surface waters that are above the
freezing point in winter. The inner AP shelf has significant topography and this roughness
may contribute to mixing (Bentley et al., 2009). At places, it is the onshore flow of the ACC
that provides the impetus to pump CDW onto the shelf, particularly along the troughs that
formed the paths of palaeo ice streams (Smith et al., 1999; Smith and Klinck, 2002; Klinck
et al., 2004).
14
Figure 1 Setting of the AP and southernmost South America showing sites discussed in text. Ocean circulation
and frontal boundaries are from Hernández-Molina et al. (2006), Orsi et al. (1995) and Naveira Garabatoa et
al. (2002) – adopted from Bentley et al. (2009).
15
2.1.2 Palaeoenvironmental records
In comparison to other Antarctic regions, the AP provides largest amount of
palaeoenvironmental records predominantly due to the large ice-free areas, each of which
present different information on the patterns, processes and mechanisms of the (Late)
Holocene climate and environmental evolution.
Ice cores, considered as one of the most solid and long-term type of record, provide an
information on environmental change in the interior from which it is possible to obtain past
atmospheric temperatures, average regional sea-ice extent, precipitation and changes in
atmospheric composition (Bentley et al., 2009). Aristarain et al. (1986) were among the firsts
who has carried out an analysis of deuterium isotope (δD) content of Dallinger Dome (James
Ross Island – JRI, AP) within a period AD 1850–1980. This has been followed by Peel and
Mulvaney (1992), who analysed δD, δ18
O and Cl–
content in Dolleman Island ice core (AD
1795–1986), to infer changes in weather and air mass patterns within the Weddell Sea zone.
Furthermore, Thompson et al. (1994) examined ice samples (oxygen isotope ratios, dust
content, chemical species, net accumulation) from Dyer Plateau, covering the period AD
1510–1989. The most recent study, performed in previous years by Mulvaney et al. (2012),
reveals nowadays the most accurate and comprehensive dataset of climate conditions of the
entire Holocene above the JRI and north-eastern AP, respectively.
Marine sediments represent another valuable environmental archive, although extracting
such records requires rather expensive and more sophisticated apparatus. Marine sediment
from various environments, such as abyssal plains, offshore continental shelfs and fjords
provide information about variation in sedimentation rate during previously triggered
processes, such as deglaciation (e.g.; Heroy and Anderson, 2005; Evans et al., 2005).
Furthermore, it is also possible to obtain information about palaeoceanographic changes –
specifically surface water productivity (Sjunneskog and Taylor, 2002), the proximity and
stability of ice shelves and glaciers (Pudsey and Evans, 2001; Brachfeld et al., 2003), extent
of sea-ice (Gersonde et al., 2003), the influx of meltwater and terrigenous sediments
(Domack et al., 1994), and clues to the behaviour of ocean currents such as the ACC (Howe
and Pudsey, 1999).
Another type of archive – lake sediments, are also frequently used for palaeo
reconstructions (onto the AP). This sort of sediment provides not only a record of
deglaciation, but also a record of changing biological productivity (linked to temperature),
16
changes in ecology and species assemblages, and lake ice cover (Hodgson and Smol, 2008).
Near-coastal lakes or lagoons can also be used as isolation basins to determine relative sealevel
change (Bentley et al., 2005a). Epishelf lakes can provide proxy records of ice shelf
presence and absence (Roberts et al., 2008).
Glaciomorphological records from land surface and continental shelf provide
supplementary information on past glacier extent and thickness (land-terminating glaciers),
and the timing of glaciation and/or deglaciation in individual areas (Sugden et al., 2006). A
variety of onshore geomorphological features have been used to determine AP ice sheet
extent, including till stratigraphy, moraines, striations, trimlines and erratics analyses (e.g.
Clapperton and Sugden, 1982; Rabassa, 1983; Bentley et al., 2006). Marine glacial
geomorphology, as an offshore-focused discipline, uses widespread technique of titled swath
bathymetry (i.e. sonar mapping) to obtain closer look onto the character of landforms on the
continental shelf seabed (Wellner et al., 2006).
2.2 Svalbard Archipelago
2.2.1 Physical setting and geographical description
Svalbard is an Arctic archipelago located north of the Norway between Greenland and Franz
Josef Land in the Arctic ocean. The archipelago is composed of nine main islands
(Spitsbergen – formerly West Spitsbergen, North East Land, Edge Island, Barents Island,
Prins Karls Foreland, Kvit Island/Gilles Land, Kong Karls Land/Wiche Islands, Bjørn/Bear
Island, Hopen), located between 10°–35° E and 74°–81° N, while the Spitsbergen is the
largest one and the only one with permanent settlements (Moholdt et al., 2010).
The Central Svalbard, including Billefjorden region consist mainly of Carboniferous and
Permian sedimentary rocks, which are divided into three groups. The Billefjorden Group
consisted largely of terrestrial deposits, which formed under the humid sub-tropical climate
during the Early Carboniferous. The Gipsdalen Group represents a mixture of siliciclastic
and calcareous sediments, and followingly Tempelfjorden Group (Kapp Starostin
Formation) is represented by sediments, which were deposited during a transgression
(Harland et al., 1997).
From a climate perspective, the general large-scale air currents are determined by the lowpressure
area near Iceland and the relatively high-pressure area over Greenland and the
17
Arctic Ocean. This resulted in transport of mild air from lower latitudes towards Svalbard.
Further north, the circulation is mostly anticyclonic with prevailing easterly or northeasterly
winds. Large temperature differences occur between the two air masses originating from the
southwest and northeast, causing extreme fluctuations in weather and temperature. The
greatest variations occur in winter when the contrasts between the two air masses are most
marked. When snow accumulation is measured on the glaciers in spring, traces of mild
periods can be often seen as ice layers in the snow pits. However, snow may fall at any time
during the summer months. On the western coast of Spitsbergen, the average annual
temperature is about –6 °C, and it is slightly and more continental further inland. The average
temperature on the western coast in the warmest month, July, is about 5–6 °C, while in the
coldest period, January – March, it is about –15 °C. Precipitation is about 400 mm annually
on the western coast of Svalbard and half of it in central inland areas. Ocean currents and
general air circulation explain the relatively mild climate. Part of the warm Norwegian
Current, a branch of the Gulf Stream, flows into the Barents Sea and part along the western
coast of Spitsbergen (West Spitsbergen Current – WSC), where it creates the northernmost
area of open water in the Arctic in winter (Førland et al., 2011).
The total glaciated area of Svalbard is 34,560 km2
, representing about 6 % of the
worldwide glacier cover outside of Greenland and Antarctica. Spitsbergen is the most alpine
island, having small cirque glaciers as well as extensive ice fields and valley glaciers. The
eastern islands, facing the Barents Sea, have less relief and are dominated by low-elevation
ice caps. Most glaciers and ice caps are considered to be polythermal (e.g., Bjornsson et al.,
1996), and 60 % of the glaciated areas drain into tidewater glaciers (Blaszczyk et al., 2009).
Glacier dynamics are typically slow with velocities < 10 m.yr-1
(Hagen et al., 2003a), but
surge events has been observed over most of Svalbard (Hamilton and Dowdeswell, 1996),
also in recent years (Sund et al., 2009).
18
Figure 2 A: Overview map of the Arctic, highlighting the polar (Arctic) Circle, the Nordic seas sector of the
Arctic Atlantic as modelled by Zhang et al. (2016), poleward circulation patterns, modified after Larsen et al.
(2013) and September 2015 sea-ice coverage. B: High Arctic Svalbard Archipelago, indicating East and West
Spitsbergen currents after Winkelmann and Knies (2005). Modified from van der Bilt et al. (2018).
2.2.2 Palaeoenvironmental records
In the Svalbard Archipelago, there are numerous extensive palaeoenvironmental records,
which provide environmental observations within the context of the last millennium. Unlike
the Antarctic Peninsula, a lot more paleo-research comes from Svalbard Archipelago, which
entails more-detailed view on the palaeoenvironmental evolution. This is associated with the
fact, that author will only present a review of a limited number of published studies,
considered as the most relevant to this thesis.
Major part of the previously published studies are the reconstructions focused on the
ocean-atmosphere interactions based on marine sedimentary cores. Specifically, Müller et
19
al. (2012) infer changes in the sea-ice conditions through a IP25 proxy in the Fram Strait for
the last 8550 yrs b2k (sea-ice cover during the last 650 yrs b2k). Spielhagen et al. (2011)
presents a multidecadal-scale record of ocean temperature variations during the past 2050
yrs b2k, derived from marine sediments off western Svalbard (79 °N). Sarnthein et al. (2003)
shows a sea surface temperature reconstruction based on a planktonic foraminifera
biostratigraphy for the last 10,750 yrs b2k in the eastern Svalbard. Dylmer et al. (2013) reveal
composite record based on three marine sedimentary deposits, in order to investigate changes
in the poleward flow of Atlantic waters (mostly within a Barents Sea region) over the last
3050 yrs b2k, using a coccolith proxies.
Furthermore, some of the researches focus as well on the ice core records. Isaksson et al.
(2001) retrieved 121-metre-long ice core from the highest point of Lomonosovfonna (1250
m a.s.l.), one of the largest ice cap of the Svalbard Archipelago. This record provided climate
record for the period AD 1920–1997 using numerous parameters, such as δ18
O or melt index.
Followingly, this was supplemented by Isaksson et al. (2005), who presented two ice core
records from Austfonna and Lomonosovfonna for the period since 600 yrs b2k. However,
ice cores particularly have been studied also by several other researches (not only in the
presented regions), especially since the early 1980’s (e.g.; Matsuoka et al., 1997).
Palaeolimnological research on Svalbard is also comparatively plentiful. Kaufman et al.
(2009) synthesize 2000 years of proxy data from lakes above 60 °N latitude with
complementary ice core and tree ring records, in order to create a palaeoclimate
reconstruction for the Arctic with a 10-year resolution. D’Andrea et al. (2012) present a subdecadal-
to multidecadal-scale record of summer temperature for the past 1850 yrs b2k from
lake sediments of Kongressvatnet in the western Svalbard. Van der Bilt et al.
(2015) study sediment from glacier-fed lake in the north-western Svalbard, in order to
reconstruct glacier and local climate variability. Not only temperature, but also hydroclimate
(precipitation) alters have been widely studied, e.g., Balascio et al. (2018) analyse δD
composition of leaf wax compounds (n-alkanes; C25-C31) in the lake sediment from northwestern
Svalbard to carry out precipitation chronology for the last 12,950 yrs b2k. Another
lake sediment from north-western Svalbard, which covers the entire Holocene environmental
history, is analysed by De Wet et al. (2018) using multi-proxy approach. History of the last
circa 5050 yrs b2k within the same region is also delivered by Gjerde et al. (2018).
20
Another approach for reconstruction – the study of glacial landforms on Svalbard is also
widely applied. Specifically, Rachlewicz et al. (2007) focus on a “post” Little Ice Age (LIA)
glacier retreat rates in the Billefjorden region, Central Svalbard. Subsequently, Phillips et al.
(2017) use 10
Be dating of two glacier moraines in south Svalbard. However, there are also
some less often methods, such as lichenometry of stabilized moraines (Werner et al., 1993).
To sum up, it is essential to mention, that presented studies are only a small fraction of what
has been studied overall in Svalbard.
21
3 SETTINGS
To investigate palaeoenvironmental evolution in polar areas two coring sites, far apart from
each other, were chosen to accomplish this task. First core retrieval was performed in Lake
Anonima (Lago Anónima) on Vega Island along the north-eastern coast of the AP. Second
core was retrieved from Lake Mathiesondalen 3, situated in Mathiesondalen valley,
Billefjorden region of Central Svalbard. Both sites withstand different physical-geographical
conditions (see 2.1.1, 2.2.1) as well as they have been subjected to different Holocene
development (see 2.1.2, 2.2.2).
3.1 Lake Anonima, Vega Island, north-eastern Antarctic Peninsula
Lake Anonima (63°49'20.34"S, 57°19'29.61"W; 24 m a.s.l.) is located in the northern part
of Vega Island, ~9 km east of the Ulu Peninsula and ~3 km north of the Terrapin Hill on
JRI. The lake is ~60 m long and ~40 m wide, with surface area ~2000 m2
. Bathymetric
measurements proved to be a shallow lake with two subbasins of maximum depth 4.1 m
(NE) and 4.6 m (SW), separated by shallow ridge. Lake is situated in an outwash plain
(proglacial zone) of the local glacier Bahía del Diablo (Marinsek and Ermolin, 2015),
whereas water depth at a core retrieval place was 3.8 m with a lake floor being slightly
sloped.
Local bedrock is formed by Cretaceous sediments of the Marambio Group, while
surrounding area of elevated volcanic mesas by Neogene back-arc volcanic rocks (basalts,
hyaloclastite breccias) of the James Ross Island Volcanic Group (JRIVG; Smellie et al.,
2008, 2013). Cretaceous strata consist of a set of fossiliferous marine shales, sandstones and
conglomerates of Santa Marta, Snow Hill Island and López de Bertodano formations
(Roberts et al., 2014). Proglacial valley floor is filled mainly with Holocene till with
predominant local volcanic rocks and in a smaller proportion by Cretaceous sediments (Zale
and Karlén, 1989). Along the northern coast of Bahía del Diablo, deltaic and marine deposits
produced by marine transgressions after the early Holocene deglaciation are interbedded
with till deposits. In the southern part of the valley, tills were partially reworked and are
covered by glaciofluvial sediment (Ermolin et al., 2002).
22
Figure 3 (a) and (c): Proglacial zone (outwash plain) of Bahía del Diablo; (b): Lake Anonima in the northern
part of Vega Island (AP) with Neogene volcanic mesas of Cape Well-met (right) and Sandwich Bluff (left)
formation in the background; (d): Core retrieval using flap gouge auger (all photographed by Jan Kavan).
The climate in the James Ross Archipelago is semi-arid with annual snowfall 200–500
mm.yr-1
w.e. (Hrbáček et al., 2017). Summer season is relatively short (December–February)
with mean annual temperature –7.0 °C (J.G. Mendel Station, period 2005–2010; Ambrožová
and Láska, 2016). Lake collects surface runoff and simultaneously is connected to a water
input from streams and snowmelt water with a sediment load that has formed a small delta
on the southwest coast of the lake (Chaparro et al., 2017). Despite the fact that the lake basin
is by Chaparro et al. (2017) considered as endorheic, the ice in till plain damming the lake
seasonally melts, causing fluctuations in a water level (~1 m), as well as in underground
drainage into surface streams and nearby lake systems (Jan Kavan, personal
communication). On account of that, lake is classified as through-flow here. During winter,
the lake surface is frozen and snow accumulates on the surface. With increasing temperatures
in late spring, the snow starts to melt first, followed by thawing of the ice layer. When the
summer temperature is high enough, active layer develops on the permafrost, putting the
complete water cycle into operation (Chaparro et al., 2017).
23
Local flora is represented only by bryophytes and lichens and its distribution is usually
limited due to the deficiency in liquid water (Robinson et al., 2003). They are located mainly
in close neighbourhood of the lakes or streams. On the other hand, the micro-flora, mostly
composed of cyanobacteria, green algae and diatoms is well developed in freshwater
ecosystems such as seepages, lakes and streams that are usually supplied by melting water
from retreating glaciers and snowfields (Engel et al., 2012). In lakes Boeckella poppei
(Mrazek) and Branchinecta gainii (Daday) are present (Nedbalová et al., 2017), however the
invertebrate fauna has been so far poorly studied. There is no penguin colony in the study
area, but several seabirds frequently visit the island. The nearest Adélie penguin colony is
located on Devil Island, ~3.5 km north-east of Lake Anonima. The human presence is very
limited, only summer campaigns have been regularly present.
Table 1 Physical-chemical properties of the water in Lake Anonima registered on 16 Jan. 2013.
Temp
(°C)
pH
field
conductivity field
(µS.cm−1
)
Acid neutralization
capacity 4.5 (mmol/l)
Na+
(mg/l)
K+
(mg/l)
Ca2+
(mg/l)
SO4
2−
(mg/l)
11.40 8.9 156 0.6 15.1 0.86 5.13 4.06
Cl−
TN TP DOC F−
NH4−N NO3−N NO2−N
24.3 0.32 0.0162 1.65 0.028 0.037 0.0107 0.001
3.2 Lake Mathiesondalen 3, Mathiesondalen, Central Svalbard
Lake Mathiesondalen 3 (78°33'42.60"N, 16°35'28.90"E; 34 m a.s.l.) is situated in the
Mathieson Valley on the eastern coast of the Billefjorden in Central Svalbard. Studied site
lies in the western part of the Sassen–Bünsow Land nasjonalpark. Lithostratigraphically,
sequences in the valley (surrounding slopes) are correlated with Minkinfjellet,
Wordiekammen and Gipshuken formations (Carboniferous–Permian), all of them belonging
to Campbellryggen Subgroup, Gipsdalen Group and Dickson Land Subgroup (Dallmann et
al., 2004). Formations consist predominantly of carbonate and evaporate rocks – dolomite,
limestone and gypsum. Late Permian high-elevated peaks belong to Kapp Starostin
Formation (Tempelfjorden Group), dominated by chert, siliceous shale, sandstone and
limestone (Dallmann et al., 2004). The outer part of the valley, where the Lake
Mathiesondalen 3 is located, is covered by Holocene raised beach (uplifted marine terrace)
sediments in the form of terraces and beach ridges (Feyling-Hanssen, 1955). The valley floor
itself is occupied by a braided river system, and small tributary streams which have eroded
gullies in unconsolidated sediments (Salvigsen et al., 1983).
24
Figure 4 (a): Search for the core’s retrieval point in the Lake Mathiesondalen 3; (b): Field coring preparation
(surrounding/background slopes formed mainly by dolomite, limestone and gypsum); (c): Gravity corer used
for extraction (all photographed on 21/08/2015 by Matěj Roman and Jan Kavan).
The lake is ~115 m long and ~55 m wide, with surface area ~5700 m2
. Geomorphologically
is this water body placed on the surface of uplifted marine terrace (34 m a.s.l.), while the
lake catchment is supplied mostly by meltwater from Fairweatherbreen and other landterminating
glaciers. The local climate in the fjord head of Billefjorden is more continental,
meaning mean annual temperature oscillates around –6 °C, while the winter season is
distinctive by temperatures around –15 °C (Gibas et al., 2005; Láska et al., 2012). Mean
annual precipitations do not exceed 500 mm, though in central Svalbard the total sum is
around 200 mm. In general, climate is highly influenced by West Spitsbergen Current
(WSC), making the Fram Strait (west of Svalbard) absent of sea-ice (Aagaard et al., 1987),
although in inner fjords (Billefjorden) the sea-ice is rather present during the winter
(December – May) season (Láska et al., 2012).
Terrestrial biota of Billefjorden is represented mainly by some megafauna species, such
as Ursus maritimus, Rangifer tarandus platyrhynchus as well as by some birds –
25
Plectrophenax nivalis, or Lagopus muta hyperborea. The lake water fauna consists mainly
of zoobenthos Chironomidae larvas, however particularly in this lake none of the diatom,
nor Cladocera species are present (Daniel Vondrák, personal communication). The nearest
permanent human settlement is ~11.7 km (north-west) far old mining Russian town
Pyramiden.
26
4 MATERIALS AND METHODS
4.1 Cores retrieval, sediment sampling and description
The drilling of the Lake Anonima sedimentary core was undertaken during the Antarctic
summer season 2012/13 (11/02/2013) by Marcos A. E. Chaparro (CONICET, Argentina),
Jan Kavan (Masaryk University, Czechia) and other expedition members. The 69-cm long
sediment core was extracted from Lake Anonima (referred to as AN13) using the flap gouge
auger. Subsequently the core was enwrapped (to prevent moisture contamination and
microbial overgrowth) in a plastic tube and after transport was stored at a room temperature
for two and half years. Prior to analyses, sediment was photographed at high-resolution,
sliced in one-centimetre steps and stored. The order of the implemented analyses, was
determined primarily by their destructiveness. Due to non-freeze storage of the sediment
samples, nitrogen content measurement was not possible.
Core from the Lake Mathiesondalen 3 (referred to as M3-15) was retrieved during the
2015 (21/08/2015) summer campaign by Matěj Roman and other members of the Polar
Ecology Course 2015, organized by University of South Bohemia (Czechia). After a
retrieval using a flap gouger auger, the 23.5-cm long sediment core was sliced in halfcentimetre
steps and stored. In both cases, compaction effect was not considered due to
uncertainty associated with degree of compaction (auger type, mass densification etc.).
4.2 Radiocarbon dating
Radiocarbon (14
C) dating of nine selected samples (Tab. 2, 3; terrestrial macrofossil remains
or, in the absence thereof, bulk sediment) was performed at Poznan Radiocarbon Laboratory
(Poznańskie Laboratorium Radiowęglowe) in Poland by means of Accelerated Mass
Spectrometry (AMS). In general, obtained 14
C ages are related to BP scale (“before
present”), conventionally referring to AD 1950. It is due to major anthropogenic influence
notable during the 2nd
half of the 20th
century (burning fossil fuels – affecting the so called
“Suess effect”; Suess, 1955 and nuclear explosions and testing; Nydal and Lövseth, 1983),
causing significant changes of radiocarbon reservoir in the atmosphere and the oceans
(Bradley, 2015). Nevertheless, all results (proxy data) in this thesis related to “age data” are
presented on a b2k (before AD 2000) time-scale with established present AD 2000.
27
Conventional radiocarbon (14
C) ages had to be calibrated so that they can be placed onto
a calendar year timescale and to eliminate differencies associated with rather inhomogeneous
spatial and temporal radiocarbon activity in the troposphere. However, using more accurate
tropical years instead of calendar ones would be more appropriate, because one tropical year
consists of 365,242192129 days, which in general results in less accurate calibration.
However, this makes not an important difference for the last few millennia of the
sedimentary records present here. Also, due to a different geographical position of sampling
sites, different calibration data sets were chosen. SHCal13 (Southern Hemisphere; Hogg et
al., 2013) curve based on dendrochronological records of terrestrial material together with
marine sediments, was applied to sedimentary core from Antarctica, whereas for a lacustrine
record from Svalbard a “global” IntCal13 data set (Reimer et al., 2013), based on a global
calibration curve, was applied. The latter is composed of dendrochronological record, and
for a period >13.9 ka cal. yrs b2k of marine (Cariaco Basin in Central Atlantic) and lacustrine
(varve chronology from Lake Suigetsu in Japan) sediments, which in addition are compared
with coral samples dated by uranium-series disequilibrium dating. Negative 14
C ages (postbomb
dates) are calibrated with the post-bomb curve (Hua et al., 2013) available for
northern, as well as for southern hemisphere. In case of Lake Anonima (Vega Island,
Antarctica) lacustrine samples, SH Zone 1–2 region data set was applied for the post-bomb
age calculation.
The age-depth modeling of calibrated ages was processed in the R-code package Clam
ver. 2.2 (Blaauw et al., 2010), using linear interpolation (Anonima) and smooth spline
(Mathiesondalen 3) method with 10,000 iterations. Calibrated best-age was processed using
weighted averages (Anonima), or mid-points (Mathiesondalen 3) of the age-depth model
derived ages. To obtain valid chronostratigraphic sequence a smoothing (0.54) was applied
in case of Mathiesondalen 3 record. Since 14
C ages are related to the appropriate depth
interval (e.g., sample AN13_5-6 equals to the sediment depth between 5–6 cm), a halfcentimetre
(Anonima) and quarter-centimetre (Mathiesondalen 3) correction had to be made,
so each (14
C) age corresponds to mid-point values, not interval ones. This step was done
directly in R code (for best-age as well as for 95 % – Anonima and 99 % – Mathiesondalen
3 confidence ranges). Subsequently, a sedimentation rate for each centimetre (yr.1 cm-1
; age
equals to mid-centimetre value) and year (mm.yr-1
) was calculated, and due to linear time
scale usage, some depth values of corresponding ages were recalculated to obtain more
realistic age-depth model.
28
4.3 210
Pb and 137
Cs (short-lived radionuclide) dating
For a recent chronology and sedimentation trend changes, the upmost 9 (Anonima) and 3.5
(Mathiesondalen 3) centimetres of the core were analysed by gamma spectroscopy for the
specific activity of isotopes 210
Pb (t1/2 = 22.3 years), 137
Cs (t1/2 = 30.1 years) and 226
Ra (t1/2 =
1600 years). The measurements were performed in the Radiometric laboratory of the
Institute of Geochemistry, Mineralogy and Mineral Resources, Charles University in Prague
by Viktor Goliáš on SILAR gamma-ray spectrometry analyser. Assumption of non-uniform
sedimentation led to application of CRS (constant rate of 210
Pb supply) model (Appleby et
al., 1979), while chronology was supplemented and supported by comparison of 137
Cs
specific activity with 137
Cs fallout curve (decline since 1963; Last and Smol, 2001).
4.4 Signal of magnetic susceptibility
Generally, low field or initial magnetic susceptibility (κ) is a measure of how easily a
material can be magnetized (Thompson and Oldfield, 1986). κ measurement was performed
at the Institute of Geophysics of the Czech Academy of Sciences in Prague on MFK1-FB
Kappabridge instrument in three measuring frequencies (976 Hz, 3 904 Hz and 15 616 Hz)
with induced magnetic field 200 mA.m-1
. To obtain mass-specific magnetic susceptibility
(χ) data, the weight of each sample together with predefined volume (~10 cm3
) was
determined to three decimal values. Thus, calculation of χ could have been then expressed
as follows:
χ =
(𝑀𝑀÷𝐻𝐻)
𝜌𝜌
, where
M stands for acquired magnetization per unit volume, and H represents (uniform) magnetic
field. In SI units, both M and H are measured in A/m, so κ is dimensionless. Hence, to obtain
the χ, κ (or equivalent M/H) should be divided by sample density ρ (Evans and Heller, 2003).
Due to a H2O being classified in terms of magnetism and on the basis of its physical
properties as a diamagnetic matter (one of the strongest; Evans and Heller, 2003), it slightly
influences magnetic signal of remaining (paramagnetic and ferromagnetic) matter and
therefore causes final data distortion. Nonetheless, all Anonima samples were measured in
their original wet conditions with a limited amount of water content (only apparent
moisture). However, by preserving the assumption of the same remaining water content
across the entire core after three years of storage, the water content may be neglected. On
29
the other side samples from Mathiesondalen 3 were measured in their natural “soaking wet”
conditions as well as in their dried form. Presented χ data are related to appropriate
frequencies, unless stated otherwise.
In order to get information of superparamagnetic (SP) grains shares in studied cores,
frequency-dependent χ (χFD), introduced by Dearing et al. (1996), was calculated according
to the following relationship:
χ𝐹𝐹𝐹𝐹 = 100 ×
𝐾𝐾𝐿𝐿𝐿𝐿−𝐾𝐾 𝐻𝐻𝐻𝐻
𝐾𝐾𝐿𝐿𝐿𝐿
, where
KLF stands for 1st
frequency (976 Hz) and KHF represents 3rd
frequency (15,616 Hz). In other
words, χFD values basically equal to percentage loss of χ and indirectly refer to ultrafine grain
content (with diameter 0–0.03 μm) in the sample.
4.5 Grain-size distribution
This method was applied to determine grain-size distribution of all samples from Anonima
core, and of selected ones from Mathiesondalen 3 core. The measurements were performed
at the Institute of Geology of the Czech Academy of Sciences in Prague on CILAS Particle
Size Analyser 1190L. Prior analysis, all samples were pre-treated in various steps to separate
foreign matter affecting the grain-size spectrum of siliciclastic constituent. Foreign matter is
in general composed of authigenic minerals such as carbonates and diatom valves (biogenic
silica) with diameter range 5–200 µm (Round et al., 1990) and their fragments. The
proportion of diatoms may vary largely, being even up to 50 % of the bulk sediment, which
would be the source material for diatomite (Kadey, 1983). All samples (2–3 g) were thus
mixed with 10% KOH solution in 120 ml acid-proof plastic tubes and then placed in an 80
°C water bath for 10 minutes. Subsequently, samples were centrifuged in Rotofix 32A
instrument for 6 minutes at 3500 rpm to separate the dissolved matter (suspension).
Nevertheless, grain-size distribution is also influenced by present organic matter. To
avoid damaging grains during pretreatment, wet oxidation was here preferred instead of
combustion organic matter removal. In general, widely used oxidants are potassium
dichromate (Walkley and Black, 1934) and hydrogen peroxide (H2O2; Schumacher, 2002;
Allen and Thornley, 2004). In this research, 30% H2O2 was used for organic matter
oxidization. Hydrogen peroxide was continually added to the sample until sample frothing
ceased. As in the previous case, the samples were heated to 80 °C during hydrogen peroxide
30
addition to increase the speed and completeness of hydrogen peroxide digestion. After the
active reaction ended, the samples were washed with distilled water by centrifugation at
3500 rpm for six minutes and the solution was decanted. As before, all chemical procedures
were carried out in high-temperature-proof (> 100 °C) and acid proof 120 ml centrifuge
tubes. On the other hand, previous carbon and sulphur content analyses revealed the presence
of carbonates below the level of detection, thus no carbonated matter removal in Anonima
sediment (grain-size in Lake Mathiesondalen 3 not assured) by 10% HCl (Vaasma, 2008)
was required.
The measurement itself proceeded along with the ultrasound particle dispersion on the
aforementioned apparatus, with a measurable range of 0.04–2500 μm. The measurement
time was set to 35 seconds, which did not include rinsing and ultrasound triggering.
Volumetric grain-size parameters (mean grain-size, clay, silt and sand content, sand/silt
ratio, mode, sorting, skewness, kurtosis, polymodality) were calculated from the raw grainsize
data in the GRADISTAT software, v 8.0 (Blott and Pye, 2001), and are presented with
graphical method (in μm) of Folk and Ward (1957).
4.6 Geochemical composition
4.6.1 Lithophile elements determination by means of X-ray fluorescence (XRF)
XRF is a nondestructive method and produces (in case of scanning approach) high-resolution
records of elemental variations (Rothwell et al., 2006; Böning et al., 2007). The elemental
data are measured as peak areas, i.e. the area under the peak of XRF spectrum, and are then
transformed to counts per unit time per unit area. This provides a relative measure of
concentration changes in the sediment. Notably also changes in sediment density, water
content, organic matter and porosity can affect XRF core measurements. According to
Moseley’s law, heavier elements emit relatively high fluorescence intensities with more
scatter and are therefore more sensitive to sample heterogeneities and mineralogical effects
such as X-ray penetration depth (Weltje and Tjallingii, 2008). The sediment pore water also
absorbs X-ray radiation, which leads to weaker fluorescence energies (Hennekam and
Lange, 2012).
Main lithophile element content through the cores was determined using the Innov-X
DELTA PREMIUM hand analyser at the Department of Geological Sciences, Masaryk
University in Brno. As for grain-size distribution, samples were measured with remaining
31
water content (Anonima, Mathiesondalen 3), or without it after drying (Mathiesondalen 3).
Every sample was measured three times in Geochem/Vanad mode, whereas final data were
arithmetically averaged. It should be noted, that during XRF measurement all samples were
stored in zip locks of the same batch meaning possible influence of storing material
elemental composition during an analysis. However, by preserving the assumption of the
same elemental composition of each plastic bag, such influence may be neglected. Besides,
the measurement of empty plastic bags confirmed that selected elements contents are several
orders lower than for sediment matter itself. Furthermore, to avoid heterogeneity and
variable water influence, ratios over content (ppm) values were preferred in results and
interpretation.
4.6.2 Carbon and sulphur content
Determination of total organic (TOC), inorganic (TIC) and total carbon (TC), together with
total sulphur (TS) was performed by a professional laboratory analyst at the Czech
Geological Survey (Brno branch) with ELTRA Metalyt-CS-1000S analyser equipped with
infrared detection. All laboratory measurements were conducted under the strict accredited
procedure (SOP-HC). To achieve highest accuracy and representativeness of the results, all
samples had been homogenized before measurement. They were dried and mill grounded,
so only bulk (homogenized) material could be extracted. After calibration and “blank test”,
when steadiness of the analyser is examined, TC and TS content of all primary sample
weights were measured, using thermal combustion (1350 °C). The exact amount of material
loss was measured via infrared detection. Secondary weight samples were measured to
determine the amount of TIC using H3PO4 (1:1). Exhibited CO2, attributed to TIC was then
detected by detector, whereas TOC component was calculated by subtracting TIC from TC.
4.7 Diatom biostratigraphy
Diatom samples were oven-dried for 24 hours at 50 °C and only 1 g of dry sediment was
used for further sample preparation. Samples were later prepared following the method
described in van der Werff (1955). Small parts of the samples were cleaned by adding 37%
H2O2 and heating to 80 °C for about one hour, followed by addition of KMnO4. After
digestion and centrifugation (3 × 10 minutes at 3700 × g), the material was diluted with
distilled water to avoid excessive concentrations of diatom valves and to unify the sample
32
volumes. Then 50 μl of every sample was mixed with 50 μl of sonicated (10 min)
microsphere solution (6792.4 MS/μl) and 400 μl distilled water and mounted in Naphrax®.
Samples and slides are stored at the Department of Ecology, Charles University in Prague,
Czech Republic.
In the Anonima core, approximately in every second sampled layer 400 valves (due to
low abundance only 50 valves in layers 33–51) were enumerated on random transects at
1,000 x magnification under oil immersion using an Olympus BX43 microscope, equipped
with Nomarski Contrast and Olympus DP27 camera using the cellSence Entry Imaging
Software. Terminology is based mainly on Hendey (1964), Ross et al. (1979) and Round et
al. (1990). For taxonomical identification of Antarctic species, Iconographia Diatomologica
vol. 24 written by Zidarova et al. (2016) was used as a main source of information. For
several species, identification up to species level was not possible due to their unclear
taxonomic situation and further morphological and taxonomical research will be necessary.
4.8 Statistical data analyses
Principal component analysis (PCA) was applied to summarise graphically the major
variation patterns of geochemical, petrophysical and selected biological data. Prior to PCA
the suitability of correlation matrix was tested by Kaiser-Meyer-Olkin (KMO) test, while
minimum data contribution in each variable was designated to 80 % except biostratigraphy
data, where diatom abundance was analysed in every second sample. Delivered correlations
(Pearson r) are significant at the 0.01 level unless stated otherwise. Biological data (diatom
species) are expressed as relative abundances (% of total diatom valves per sample). Diatom
characteristics – number of valves per gram of sediment, species richness, Shannon diversity
(Shannon and Weaver, 1949) and Evenness (Pielou, 1966) – were calculated according to
cited publications. Significant diatom clusters were identified and divided using
stratigraphically constrained cluster analysis (CONISS) with significant division based on
comparing a CONISS distance matrix with a random model based on broken stick model
(Bennett, 1996) using the rioja package in R (Juggins, 2009). All presented graphs were
produced in Grapher v. 11.
33
5 RESULTS
5.1 Lake Anonima
5.1.1 Lithology and age-depth model
Lithological profile is comparatively undifferentiated with undeveloped annual lamination
or visible stratification (Fig. 5). For an age-depth model of deposition in Lake Anonima (Fig.
5) five radiocarbon ages were obtained from the core and are presented in Tab. 2., although
the fifth and concurrently nethermost sample (66–67 cm) revealed significantly
underestimated radiocarbon age. Due to reversal age of the lowermost sample, only the
topmost four radiocarbon ages, revealing suitable chronostratigraphic sequence (goodnessof-fit
value 4.96), were included into the calculation of age-depth model with a base onset
at 2440 cal. yrs b2k (54–55 cm). Tab. 2 shows that the topmost part (5.5–0 cm) accumulated
after 1950 AD, therefore post-bomb radiocarbon age, where the amount of nuclide is
conventionally expressed as pMC entailing “percentage of modern carbon” – normalized to
100 % (Reimer et al., 2004). Short-lived isotopes (137
Cs, 210
Pb) dating proved very low
specific activities (Bq/g) in the samples, thus relationship purification between age and depth
within recent decades was not possible.
Table 2 Radiocarbon ages of dated samples from the Lake Anonima.
Sample
Laboratory
code
Depth
(cm)
Material
mg
Ca
14
C years BP
Min.–Max. age
cal. yrs b2k (± 2σ)
AN13_5-6 Poz-85803 5–6
lacustrine
moss
0.09 100.99 ± 0.39 pMC 43–44
AN13_16-17 Poz-89401 16–17
lacustrine
mud
– 555 ± 30 554–604
AN13_28-29 Poz-89403 28–29
lacustrine
mud
– 1650 ± 30 1468–1615
AN13_54-55 Poz-89402 54–55
lacustrine
mud
– 2130 ± 30 1991–2204
AN13_66-67 Poz-85804 66–67
lacustrine
mud
0.06 1260 ± 80 –
a
milligrams of carbon in the analysed sample, where the amount of carbon was < 0.1 mg
Sedimentation rate along the profile varies in range 0.13–0.97 mm.yr-1
suggesting variable
clastic input within time. Interval 2440–1560 cal. yrs b2k shows comparatively higher
sedimentation rate (0.45 mm.yr-1
) followed by long sediment input reduction between 1560–
90 cal. yrs b2k (0.13–0.21 mm.yr-1
). Near the core’s top the sedimentation rises steeply to
0.97 mm.yr-1
, which might correspond with smaller compaction of topmost part of the core.
34
Figure 5 Best estimate age-depth model (black line) with 2σ uncertainty envelope (95 % confidence interval)
based on four 14
C ages derived from the Lake Anonima. The lowermost reversal 14
C age is not included in agedepth
model calculation. Supplemented by probability distribution curves of calibrated ages, sedimentation
rate (mm.yr-1
), mass magnetic susceptibility (χ976Hz) and diatom biozonation (Bennett, 1996 model).
5.1.2 Physical proxies
Measurements of χ and χFD were performed with samples in original wet conditions with a
limited quantity of water content (merely apparent moisture). Still, by preserving the
assumption of the same remaining water content across the entire core after three years of
storage, paramagnetism and diamagnetism of the water might be neglected. The values of χ
obtained throughout the core are rather low (generally in the range of 10–9
m3
.kg-1
) showing
a limited influence of ferri- and ferromagnetic particles and dominance diamagnetic and
paramagnetic material.
Lowest part (2440–2200 cal. yrs b2k) of the core section is distinctive by comparatively
low χ vacillating around ~3 × 10–9
m3
.kg-1
, followed by emergent shift (2200–2000 cal. yrs
b2k) to increased values (~13 × 10–9
m3
.kg-1
) towards the middle section of the core (2000–
1400 cal. yrs b2k). Abrupt apex at 1400 cal. yrs b2k initiated continuously decreasing drift
35
ceasing at 550 cal. yrs b2k (~8 × 10–9
m3
.kg-1
). Time interval 550–0 cal. yrs b2k exhibited χ
variation in the range ~8–10 × 10–9
m3
.kg-1
. Calculated proxy parameter χFD revealed
apparent variability throughout the profile. Nonetheless, lowest part shows relatively higher
χFD (~5 %) with a rapid decrease at 2100 cal. yrs b2k (to ~3.5 %), corresponding to opposite
shift in χ. Time interval 2000–0 cal. yrs b2k reports continuously increasing course (~3.5 to
5 %) towards the core top.
Figure 6 Physical proxies (left to right): mass magnetic susceptibility χ (10-9
kg-1
.m3
), frequency-dependent
magnetic susceptibility χFD (%), grain-size fractions (%), mean grain-size MGS (µm), sorting degree,
skewness and sedimentation rate (mm.yr-1
) plotted against b2k time-scale.
Sediment has predominantly unimodal distribution with lowermost 15 cm (2440–2110 cal.
yrs b2k) and topmost 19 cm (780–0 cal. yrs b2k) being poorly sorted. Middle section (2110–
780 cal. yrs b2k) comprises changes of sorting in each 5–7 cm interval from moderately
sorted to poorly sorted and conversely, respectively. Degree of sorting is appositely reflected
in grain-size distribution, where sections with higher shares of sand fractions tends to be less
sorted. Anyhow last 50 cal. yrs b2k are characterised by increasing sand (~5 %) and
decreased clay (~17 %) content correlating with higher mean grain-size (MGS; ~10 %).
36
Period 50–1500 cal. yrs b2k shows relatively static clastic input with clay fraction oscillating
around 28 % (silt ~72 %) whereas sand fraction is not present. Lower part of the core has
more inferior the same fraction proportion, although at 1610 and 2070 cal. yrs b2k two steep
peaks of clay fraction (up to 58 and 49 %) are apparent. The oldest part of the profile
demonstrates diminishing curve in clay fraction down to 12 % with sand fraction increase
up to 13 %. MGS parameter variation correlates well with fine sand fraction (r = 0.93).
5.1.3 Geochemistry
Whilst none inorganic carbon was detected in the entire core, TOC and TS courses
correspond to each other with their highest relative content (maximum TOC 5.1 % and TS
0.2 %) in the bottom section (2440–2130 cal. yrs b2k). Organic carbon and sulphur then
rapidly reduces almost to zero, which is compensated in upper section (1420–260 cal. yrs
b2k), where TOC oscillates around 1 %. Fluctuations linkage with subsequent inorganic
proxies is best expressed with light elements (LE) content (r = 0.90) and Rb/Sr (r = 0.87)
plus Fe/Mn (r = 0.86) ratios. Direct relationship between the latter two is also apparent (r =
0.89), whereas biogenic silica (BSi) presence expressed circuitously by Si/Ti ratio as well
corresponds with TOC (r = 0.81) very fitly. Grain-size Zr/Ti and Al/Si proxies also bear
cognation (r = –0.86) establishing these two ratios as a solid proxy for grain-size distribution.
To distinguish main processes, material sources in the catchment area and heterogeneity and
variable water influence in the core, the author relies on ratios over ppm content of individual
elements.
37
Figure 7 Geochemical proxies (left to right): total organic carbon TOC (%), Rb/Sr, Fe/Mn, Si/Ti, Zr/Ti and
Al/Si ratios, Si, Fe, Ca, Al, Ti and Mn contents (ppm).
5.1.4 Diatom biostratigraphy (by Kateřina Kopalová and Marie Bulínová)
The complete diatom flora of the Lake Anonima core reveal 80 species in total, belonging
to 26 genera. Based on the CONISS cluster analysis the Lake Anonima core is being divided
into six diatom zones (from now on a “biozone” term is used), four of which are statistically
significant based on the broken stick model (biozones 1–4). Subdivisions within the biozone
3 are not statistically significant, but they are retained due to their correspondence with
changeable TOC content within the biozone 3 (Fig. 8). Diatom productivity mirrors patterns
in TOC, which is the highest in biozone 1, where the diatom species reach their highest
abundance 604 × 104
valves per gram of dry sediment. The lowest productivity subsists in
the biozone 2 where the observed maximum is only 10 × 104
. More intermediate and variable
is in the biozones 3 and 4 (ranging 20 × 104
–210 × 104
valves per gram of dry sediment).
Richness increases overall from the biozone 1 to the biozone 4 with the peak in the recent
material in the biozone 4, and being lowest in biozones 1 and 2. Shannon diversity is more
stable, but has the lowest values in biozone 1 ranging between 1.36 and 2.49 and the greatest
values in biozone 4, where the maximum value of 2.89 is remarked. Evenness shows a more
38
fluctuating pattern with the greatest value in biozone 2 reaching its maximum 0.93. The
diatom community as a whole is dominated by the genus Nitzschia (average share of the
genus in the core is almost 41 %) and the abundances are relatively uniform within the core.
The next most abundant taxon is Planothidium rostrolanceolatum, which occurs in share up
to ~ 41 % in the biozone 3. All other taxa contribute by smaller proportions and are more
variable within the core. Specific patterns in diatom community in individual biozone
follow.
Figure 8 Diatom stratigraphy of the Lake Anonima sedimentary core plotted against b2k time-scale. Zonation
is based on a CONISS analysis of the diatom species composition. Only diatom species with a relative
abundance >8 % are depicted. Significant group divisions are marked with red lines. Photomicrographs of
presented species are to be found in Zidarova et al. (2016). Supplemented by TOC (%) and valve
abundance/productivity (both left).
Biozone 1 (2440–2050 cal. yrs b2k) is dominated by the genus Nitzschia (46–67 %),
followed by Achnanthidium australexiguum (4–22 %), Sellaphora antarctica (4–16 %),
Stauroneis delicata (0.2–10 %, being mostly present only in the biozone 1 and subzone 3c),
and Gomphonema maritimo-antarcticum (0.7–11 %). Brachisira minor, the most abundant
species in the biozone 2, shows almost no presence in biozone 1. Diatom biozone 2 (2050–
1520 cal. yrs b2k) is different, than all other biozones characterized by very low total diatom
valve concentration (Fig. 8). In the biozone 2, only 50 valves per sample were counted in
comparison to 400 in other biozones due to extremely low valve concentration, which may
have artificially reduced the calculated species richness and increased Evenness. The diatom
39
communities present mostly (semi-)terrestrial diatom species such as Brachisira minor
showing the highest abundance in the biozone 2 (2–28 %), Hantzschia amphioxys (0.5–10
%), Luticola muticopsis (0.5–18 % appearing only in the biozone 2), and also genera
Humidophila is present. In much of biozone 2, Sellaphora gracillima is absent. The
community is still dominated by genus Nitzschia, but within this genus Nitzschia
homburgiensis comprises up to 18 % of the community in contrast to the rest of the core,
where it appears in very low abundances.
Figure 9 Diatom characteristics: (Diatom species) richness is the number of taxa in the sample, Valve
abundance presents diatom valves per one gram of dry matter (sediment), Shannon diversity summarises the
diversity of a population in which each member belongs to a unique group (species), and species Evenness
refers to homogeneity of the species. The zonation is based on the CONISS analysis of the diatom species
composition (Fig. 8).
Diatom biozone 3 (1520–220 cal. yrs b2k) is subdivided into three sub-zones (3a: 1520–980
cal. yrs b2k, 3b: 980–460 cal. yrs b2k and 3c: 460–220 cal. yrs b2k). Biozone 3 is
characterised by co-dominance of Planothidium rostrolanceolatum and the genus Nitzschia.
Planothidium rostrolanceolatum has the highest abundance in the sub-zone 3c (14–32 %).
Achnanthidium australeexiguum, which is present in sizable amount throughout the core is
40
almost absent in biozone 3a. The characteristic feature of biozone 4 (220 cal. yrs b2k –
current age) is the reduction of the abundance of Planothidium rostrolanceolatum and
increase of abundance of Sellaphora gracillima. Humidophila keiliorum is present again
after being almost absent in biozone 3.
5.1.5 PCA
First PCA (total explained variance 84.10%; KMO value 0.83 – “meritorious” sample
adequacy; Fig. 13) showed that major variability of non-biotic data together with diatom
indexes within a core is driven mainly by PC1 with 61.73 % of explained variance. PC2
(9.84 %) together with PC3 (6.75 %) and PC4 (5.78 %) are not considered as substantial due
to their low explained variance and thus low interpretable clarity. PC1 bears a strong positive
correlation with LE (r = 0.97), TOC (r = 0.94), Rb/Sr (r = 0.92), Fe/Mn (r = 0.91), Zr/Ti (r
= 0.84), Rb (r = 0.81), Si/Ti (r = 0.80), sand fraction (r = 0.78), TS (r = 0.76) and also strong
negative correlation with Fe, Sr, Ca (r = −0.98), Zn (r = −0.97), Cu (r = −0.96), χ976Hz (r =
−0.95), Mn, Al/Si (r = −0.94). PC2 has an appropriate reflection with Shannon and Evenness
(r = 0.96), richness (r = 0.92), but only fair correlation with productivity (r = 0.50).
Second PCA (KMO value 0.79 – “middling” sample adequacy; Fig. 14) suggested lower
degree of total explained variance (72.42 %) due to probable influence of increased data
heterogeneity (several diatom species) in the analysis. PC1 however explains 53.28 % of
total variance within sediment, and unlike in the previous PCA (Fig. 13), a PC2 explains
slightly more input data variability (12.64 %). Nevertheless, due to a low explained variance
(6.50 %) a PC3 is not considered to explain a substantial process role during sedimentation,
thus it is not included in interpretation. Most significant loadings (correlations) in the PC1
correspond to PC1 of the previous analysis – i.e. organic matter represented by TOC (r =
0.93) and LE (r = 0.97) majorly correlates positively with PC1. As well Rb/Sr (r = 0.91),
Fe/Mn (r = 0.90), Zr/Ti (r = 0.84) and Si/Ti (BSi r = 0.80) elemental ratios do have a
significant positive relationship, while binding between petrophysical proxies and PC1 is
still significant, but less intense (sand r = 0.79, χFD r = 0.58). Indirect link of PC1 is visible
with a few elements (Fe, Ca, Sr, Zn r = –0.97; Ti r = −0.86; Al r = −0.79), as well as with
elemental ratio Al/Si (r = −0.94) and petrophysical proxies (clay r = −0.65, χ976Hz r = −0.95).
Diatom assemblages are associated more with PC2 (Sellaphora Antarctica r = 0.63,
Sellaphora gracillima r = 0.56, Brachysira minor r = 0.54), while diatom characteristics
41
relate to PC2 as well (Shannon, Evenness r = 0.95; richness r = 0.89), except productivity,
which has littler same correlation coefficient with both components (PC1 r = 0.54, PC2 r =
0.49).
5.2 Lake Mathiesondalen 3
5.2.1 Lithology and age-depth model
Due to direct field sub-sampling of the core, lithostratigraphic description was not taken into
consideration, meaning that lithological units throughout the sediment were not later
identified and described. In addition, prevailing absence of terrestrial macrofossils led to the
analyses of bulk sediment samples. However, four radiocarbon ages supplemented by
uppermost seven samples measured for short-lived (210
Pb, 137
Cs) isotopes, revealed
comparatively robust chronological sequence (goodness-of-fit value 9.85), which covers the
last ~700 years of natural history.
Table 3 Radiocarbon ages of dated samples from Lake Mathiesondalen 3.
Sample
Lab
code
Depth
(cm)
Materiala
mg
C
14
C
years BP
Min.–Max. age
cal. yrs b2k (±
3σ)
Min. – Max.
(including reservoir
4380 ± 40 14
C BP)
M3-15_1.5-2
Poz-
96535
1.5–2 lac. mud 0.35
4330 ±
40
4837–5031 –
M3-15_8.5-9
Poz-
96536
8.5–9 lac. mud 0.6
5500 ±
40
6214–6398 935–1174
M3-15_16.5-
17
Poz-
96537
16.5–
17
lac. mud –
5080 ±
40
5739–5913 556–728
M3-15_23-
23.5
Poz-
96538
23–
23.5
lac. mud –
4840 ±
40
5476–5652 330–625
a
”lac. mud” = lacustrine mud
Table 4 Short-lived isotope activity and associated ages of the core top from Lake Mathiesondalen 3.
Sample/depth (cm)
specific activity [Bq/g] Calculated yrs b2k
(± 1σ)210
Pb (± 1σ) 137
Cs (± 1σ) 226
Ra (± 1σ)
0–0.5 0.330 ± 0.024 0.027 ± 0.008 0.141 ± 0.013 −14 ± 4.0
0.5–1 0.316 ± 0.013 0.031 ± 0.004 0.130 ± 0.007 −12 ± 3.8
1–1.5 0.300 ± 0.015 0.023 ± 0.005 0.156 ± 0.009 −5 ± 3.8
1.5–2 0.269 ± 0.007 0.028 ± 0.002 0.135 ± 0.004 −1 ± 3.5
2–2.5 0.250 ± 0.007 0.040 ± 0.003 0.133 ± 0.004 10 ± 3.7
2.5–3 0.257 ± 0.009 0.043 ± 0.003 0.138 ± 0.005 22 ± 3.9
3–3.5 0.194 ± 0.007 0.047 ± 0.003 0.126 ± 0.004 44 ± 3.6
42
Figure 10 Best estimate age-depth model (middle bold black line) with 3σ uncertainty envelope (99 %
confidence interval) based on three 14
C and seven short-lived (137
Cs, 210
Pb) isotope (inset) based ages.
Combination of such methods enabled establishment of upmost 14
C age (red star marks) being equal to -1 yrs
b2k. The 8.5–9 cm reversal 14
C age is not included in age-depth model calculation. Supplemented by
probability distribution curves of calibrated ages (3σ) and sedimentation rate (mm.yr-1
).
Based on the comparison of two independent methods (14
C vs. 210
Pb and 137
Cs) for the
material form 1–1.5 cm depth, a radiocarbon reservoir was determined at 4380 ± 40 14
C yrs
BP, whereas significantly overestimated radiocarbon age (5500 ± 40 14
C yrs BP) at 8.5–9
cm depth was not taken into account for the age-depth model build-up. Since a smooth spline
method was applied for an age modelling, sedimentation rate does not change abruptly, but
rather continuously with less pronounced shifts. Rate of material input alters within a range
of 1.85–25 mm.yr-1
, with its highest sedimentation rate (25 mm.yr-1
) being distinctive for
the lowest section (~23 cm; ~700 cal. yrs b2k). It is followed by a continual decrease (25
mm.yr-1
→ 1.85 mm.yr-1
) to 300 cal. yrs b2k (8.75 cm), and in the upper part of the
succession, the sedimentation rate steadily rises to 10 mm.yr-1
at the core’s top, which might
43
be an artefact of uneven distribution of dated samples. Handed-down uncertainty associated
with the compaction effect is, also here, on point.
5.2.2 Geochemical and physical proxies
The carbon and sulphur content proved significant changes over time. Continual increase in
TC (8.91 % → 10.59 %) towards the present is apparent, while opposite course is marked
by TS (0.13 % → 0.07 %). TOC content shows comparatively increased variability, although
between 700–380 cal. yrs b2k a recorded decrease (~5.98 % → 4.50 %), with following
(since 380 cal. yrs b2k till present; ~4.50 % → 6.23 %) opposite (and more variable)
increase, is evident. Unlike TOC, TIC shows in the time range 700–380 cal. yrs b2k a
continuous increase (2.93 % → 4.50 %) with succeeding, however more variable trend
towards the present.
Inorganic geochemistry results (ppm content and ratios) are presented in dried (d) as well
as in original wet conditions (w – in Figs. 11 and 12 titled as “+H2O samples”). However,
only few of the presented elemental contents (Fig. 11) do bear a considerable course and/or
variability. Generally, the moist samples show significantly lower contents of chemical
elements (Fig. 11) due to variable water absorption of X-ray radiation. From this perspective,
the XRF data obtained from dried samples should be preferred. For instance, Si and Al (both
d) contents decrease since ~380 cal. yrs b2k, while Fe and Ca (w+d) show rather long-term
decrease and increase, respectively. Selected elemental ratios (Fig. 11) reveal permanent
increase (e.g., Si/Ti, Zr/Ti), while other ratios reveal decrease (e.g., Rb/Sr, Fe/Mn).
However, notably, values of all ratios (with an exception of Rb/Sr) do fluctuate quite
sensitively, resulting in more wavering and less indicating trends. Al/Si ratio does not
suggest any long-term course.
Due to logistic issues, the magnetic susceptibility (χ) of the upmost 5 cm of dried samples
were not analysed, thus they are not presented in the results. However, as in the Lake
Anonima χ values obtained throughout a core are rather low (both d and w – generally in the
range of 10–8
m3
.kg-1
) showing a limited influence of ferri- and ferromagnetic particles and
concurrently dominance of diamagnetic and paramagnetic matter. Also, all three frequencies
of χ show persistent decreasing trend since 700 cal. yrs b2k. Nevertheless, magnitude of
decrease is more intense for dried samples between the period 700–500 cal. yrs b2k. Due to
44
the overlapping χ of different frequencies (w and d), the χFD parameter displays negative
values (w and d).
Relationships between variables of interest have been revealed using correlation counts,
however for some tight connections, the independence of both variables was violated (i.e. r
= 0.97 between Rb/Sr and Rb), which have resulted in exclusion from the analysis. Separate
measurements of dry samples revealed that strongest positive coherence is present between
Rb and Zn (r = 0.98), Si and Al (r = 0.94), and Rb/Sr and Zn (r = 0.92). There is also strong
positive correlation of Rb/Sr with Fe/Mn (r = 0.81), χ3904Hz (r = 0.88) as well as with χ15616Hz
(r = 0.86), however very weak with χ976Hz (r = 0.46). Separate Rb content is linked well with
χ3904Hz (r = 0.81) and Fe content (r = 0.86). Fe correlates quite fitly with Zn (r = 0.87), χ3904Hz
(r = 0.86) and χ15616Hz (r = 0.85). Also, Si and K relationship corresponds rather well (r =
0.85). On the other side, the strongest indirect correlation is mostly associated with TC
behaviour. Specifically, strongest negative correspondence of TC is with Rb (r = –0.94),
Rb/Sr and Zn (both r = – 0.91), Fe/Mn (r = –0.87), Ni (r = –0.84), Cu (r = –0.79), K (r = –
0.77), Zr (r = –0.75), Cr and χ3904Hz (both r = –0.74). This is supplemented by Ca and K (r =
–0.87), and indirect connection of Sr with Fe/Mn (r = –0.75).
Analyses of moist samples also revealed certain substantial linkages among variables.
Particularly, positive coherence is tightest between Zn content and K and Rb (both r = 0.95),
respectively. As in the context of dried samples, Rb content behaviour is linked quite tightly
with K (r = 0.94), Zr (r = 0.90) and Fe (r = 0.89). Furthermore, Al corresponds highly with
Ti (r = 0.93), while Si/Ti is linked with Zr/Ti (r = 0.88). Strong positive relationship is also
between Rb content and χ976Hz, χ3904Hz (both r = 0.91) and χ15616Hz (r =0.90). On the other
hand, some of the strong indirect correlations are related to TC variability throughout a
profile. TC corresponds negatively mostly with Rb (r = –0.94), Rb/Sr (r = –0.88), Zr (r = –
0.85), K (r = –0.84), Zn (r = –0.83), Fe (r = –0.80) as well as with all three χ frequencies –
χ976Hz (r = –0.84), χ3904Hz (r = –0.84), χ15616Hz (r = –0.83). LE content behaviour within dried
samples is in line with Ca (r = –0.81) and Sr (r = –0.76), while sufficient negative
relationship is also between Al content and Zr/Ti (r = –0.67), and Ca and Rb/Sr (r = –0.79).
Further correlations and linkages among variables are provided by PCA results, presented in
5.2.3.
45
Figure 11 Geochemical organic proxies (TC, TIC and TOC, TS – all %) together with selected elemental ratios
(Rb/Sr, Fe/Mn, Si/Ti, Zr/Ti and Al/Si) plotted against b2k time-scale.
Figure 12 Physical proxies (left) together with elemental content (centre to right) plotted against b2k time-
scale.
46
5.2.3 PCA
PCA was performed on all geochemical and physical data to obtain information about the
dominating patterns within a core composition. However, for more-detailed interpretation
PCA was performed on both moist and dried samples, meaning this has enabled taking a
separate H2O influence into the consideration.
Analysis of moist samples (Fig. 15) revealed, that core’s geochemical and physical
composition consists of six components in total, whereas only PC1 (40.40 %) and PC2
(17.81 %) could be considered as substantial due to their relatively high level of explained
variance. First component bears strong positive correlation with some alkali metal elements,
i.e. Rb (r = 0.96), and K (r = 0.90), as well as with transition metal elements – Zn (r = 0.90),
Zr (r = 0.87) and Fe (r = 0.91). This is supplemented by strong positive correlation with
χ976Hz (r = 0.93), χ3904Hz (r = 0.93), χ15616Hz (r = 0.91), Rb/Sr ratio (r = 0.93) and comparatively
strong relationship with TS (r = 0.69). In contrast, reduction of these variables results
conditionally in forcing of TC and Ca, which are in negative correlation (r = –0.93, –0.62)
with PC1. Second component positively correlates rather with Si (r = 0.63), Ti (r = 0.64),
Ca (r = 0.59), Al/Si (r = 0.56), but negatively with LE (r = –0.77).
The statistical procedure within dried samples (Fig. 16) showed, that core’s composition
consists of five components in total, however only the first two are conceived as
fundamental. PC1 explains 53.07 % of total variability, while PC2 explains only 12.20 %.
Positive PC1 agrees with increased TC (r = 0.97) and Ca (r = 0.79), as well as with Sr content
(r = 0.72). TIC (r = 0.47) and TOC (r = 0.56) positions fit appropriately within a twodimensional
PC1. On the contrary, negative feedback of PC1 is associated with increase of
variables group consisted of χ976Hz (r = –0.88), χ3904Hz (r = –0.94), χ15616Hz (r = –0.94), Rb (r
= –0.95), Rb/Sr (r = –0.94), TS (r = –0.71), Fe (r = –0.75), Zr (r = –0.76), Zn (r = –0.92),
Fe/Mn (r = –0.86), Ni (r = –0.86), K (r = –0.81) and Al/Si (r = –0.73). Al, Si and Ti contents
reveal comparatively substantial negative bound to PC1 (r = –0.79, –0.70, –0.40) as well as
considerable positive one to PC2 (r = 0.46, 0.51, 0.70). LE and χFD seem to be probably
significantly unrelated to either the PC1 (r = –0.31, 0.19) or PC2 (r = –0.52, 0.39).
Based on the abiotic variables changes a main phases (sections) were identified. The
earlier phase is defined between 700–380 cal. yrs b2k, while the latter one comprises period
since 380 cal. yrs b2k till present times.
47
6 DISCUSSION
6.1 Lake Anonima evolution in the Late Holocene
6.1.1 2440–2050 cal. yrs b2k: Productive warmer period
According to the age-depth model’s best estimate age (Fig. 5), Lake Anonima sedimentary
record spans the last ~2450 years (sediment base 2440 cal. yrs b2k) attaining thus
environmental and climatic history over the major part of the Late Holocene. Data variability
of individual proxies showed significant alternations in environmental, climatic and lake
conditions of the north-eastern AP region.
First period 2440–2050 cal. yrs b2k is distinctive by comparatively low χ (~3 × 10–9
m3
.kg-1
) relating to low content of (iron-bearing) paramagnetic and ferromagnetic minerals
(Evans and Heller, 2003), as well as to potentially warmer climatic phase (Maher and
Thompson, 1999). Increased χFD (~4.72 %) in the core links to higher share of SP grains with
diameter <0.3 μm (Dearing et al., 1996), however its literal origin is yet unresolved, since
necessary specifies for chemical weathering (good aeration, mild precipitation, near neutral
pH and warm climate) does not tally with the Antarctic environments (Worm, 1998). Based
on the outcomes of elemental and chemical composition of soils, Navas et al. (2008)
conclude that the (outcropping) bedrock type has the strongest influence, and thus physical
weathering is the main driving force of pedogenesis in Antarctica. However, Haus et al.
(2016) prove that driving element of chemical weathering within the studied region are
enhanced by tephra depositions from Deception Island. Due to these issues, χFD parameter,
or SP grains origin hence demands further and more-detailed investigations in the context of
extreme (Polar) environment.
Association of magnetic parameters is regardlessly conveyed with grain-size variation of
the sedimentary material deposited in the lake. Higher sand fraction (~13 %) and coarser
MGS (~11.8 μm) in the given time interval are in comparison to the following period (since
2050 cal. yrs b2k) associated with higher siliciclastic input into the lake basin, induced by
thawing of local glaciers and permafrost active layer. Increased input coincides well with
higher sedimentation rate (~0.45 mm.yr-1
) and appropriately with elemental ratios data,
utilized as grain-size (Zr/Ti, Al/Si; Kylander et al., 2013), or chemical proxies (Rb/Sr; Xu et
al., 2010). For explanation, Zr is associated with coarse silt, whereas Ti is bound to clay
48
fraction (Taboada et al., 2006), and similarly, clay is represented by Al, while sandy grains
by Si (Lopez et al., 2006).
This period almost precisely equals to biozone 1 (2440–2050 cal. yrs b2k) and is
characterized by higher abundance of diatom aquatic species – Achnathidium australexigum
(4–22 %), Stauroneis delicate (0.2–10 %) and Sellaphora antarctica (4–16 %). To define
most accurate ecological niche of the species from Lake Anonima, data were compared with
adjacent lakes on the JRI (Bart van de Vijver, personal communication). Achnathidium
australexigum and Stauroneis delicata are being mostly present in stable shallow lakes (pH
7.2–7.5) with low nutrient content and thus low conductivity, however Sellaphora antarctica
is oppositely distinctive for lakes with silt drapes and algal mats developed on the lake floor.
Species abundance is supplemented by highest productivity (~299 × 104
valves per gram of
dry sediment) within this period, but comparatively also by low species richness (~19) and
Shannon index (~2.2) suggesting a productive, but uniform (low biodiversity) environment
in the lake. Uniformity is supported by relatively low values of Evenness (~0.74) proposing
appropriately developed inequality within a relative abundance, meaning probable
dominance of certain group among all the species. TOC and TS show comparatively highest
content (max. 5.07 % and 0.91 %) in this period, which corresponds to highest diatom
productivity as well as to Fe/Mn redox proxy (O’Sullivan and Reynolds, 2008) values. PCA
1 (Fig. 13) however prove, that productivity parameter is explained almost equally by both
(PC1 61.73 %, PC2 9.84 %) components, meaning correlation between organic matter and
productivity, nor with indexes (Shannon, richness, Evenness), is not strictly clear. Due to
that, productivity could not be considered as a single-dependent parameter in this study.
PC1 mostly correlates with those variables, which are linked to climate fluctuations
processes, intending specifically that latent variable (PC1) change is indisputably associated
with alternation of climate variables (temperature, moisture). On top of that, increase of PC1
has positive feedback on organic matter content (LE, TOC, TS, Fe/Mn), as well as on Si/Ti
ratio, representing proxy for biogenic silica content (BSi). Tight correlations between Si/Ti
(BSi), sand fraction and (more less) productivity suggest a substantial proportion of biogenic
origin of siliceous material within a core. Above that, this is endorsed by negative (opposite)
relationship between these variables and Si, suggesting Si does not clearly have only
inorganic (lithogenic) origin. Linkage between χFD together with sand fraction and positive
PC1 suggests, that SP grains are climate dependent, and even though χFD is associated with
49
sand fraction, its origin is more probable due to a chemical, rather than physical weathering.
Rb/Sr close relationship with PC1 increase also establishes Rb/Sr as a proxy for chemical
weathering, rather than for chemistry changes conditioned by different lithology. Usage of
Rb/Sr proxy as a provenance indicator is quite debatable; therefore, it demands further
investigations in the context of Antarctic environment. On the other side, negative PC1
influence results in positive feedback of clay fraction, meaning the presence of colder
conditions, which is in correspondence with sediment fining. This is supported by close
correlation of this variable with Al/Si (grain-size) proxy, and opposite correlation with Zr/Ti
(grain-size) proxy. Tight link of these variables with χ976Hz suggests that susceptibility signal
is bounded to clay fraction (magnetite, maghemite).
Figure 13 Lake Anonima: PCA 1 (two-dimensional) biplot plotting PC1 (61.73 %) against PC2 (9.84 %) along
with all abiotic and selected biotic variables being colourised by data type. Note space distortion (variable
positions) due to PC2 axis adjusted range.
PC2 probably explains changes, which do not feature linkage to climate fluctuations,
entailing latent component as chemically conditioned by different lithology. Position of
diatom species close to the PC2 axis suggests, that assemblages are chemically, rather than
climate-dependent. This is especially true for diatom species typical for biozones 2
(Brachysira minor, Nitzschia homburgensiensis) and 3 (Planothidium rostrolanceolatum,
Sellaphora gracillima). However, some of the species (mostly biozone 1 – Sellaphora
50
antarctica, Stauroneis delicata, Achnanthidium australexiguum) do have noticeable positive
relationship with PC1 as well, meaning those assemblages are also linked with climate
factor.
Figure 14 Lake Anonima: PCA 2 (two-dimensional) biplot plotting PC1 (53.28 %) against PC2 (12.64 %)
along with all abiotic and selected biotic variables being colourised by data type. Numbers in brackets over the
diatom name constitutes for which biozone the diatom species is representative. Note space distortion due to
PC2 axis adjusted range.
6.1.2 2050–1520 cal. yrs b2k: Colder phase → decline in primary productivity
Subsequent zone time interval is expressed by biozone 2 (2050–1520 cal. yrs b2k), in which
selected biotic proxies report high level of diminishing (TOC ~0.12 %, TS < limit of
detection, productivity ~2.63 × 104
, in part richness ~19.2). This is supplemented by
corresponding data trends within abiotic proxies. Specifically, MGS parameter decreases
(~11.8→5.5 μm), which is in coincidence with higher clay content (~31.3 %), Al/Si values
(~0.26), and on the other side decreased sand fraction (~0.03 %) together with Zr/Ti (~0.02).
Re-dox proxy (O’Sullivan and Reynolds, 2008; Fe/Mn) potential is also relatively low
(~59.31). χFD shows lowered intensity of chemical weathering (Chen et al., 1999; ~3.74 %),
while χ suggests higher content (~13.07 × 10–9
kg-1
.m3
) of paramagnetic to ferromagnetic
matter in the sediment (Evans and Heller, 2003). Decreased biogenic silica (BSi ~6.37) is
also in agreement with biotic proxies’ trend. In sum, all of these proxies disclose colder
51
climatic phase, most probably associated with limited meltwater, or precipitation input into
the lake basin, which might have later resulted in lake level drop. Predominant (diatom)
aerophilic species (Brachysira minor, Nitzschia homburgiensis) refer to water diminishing
due to drier conditions. Brachysira minor (not found in the biozone 1) is connected mostly
to these drier and colder conditions, whereas Nitzschia homburgiensis have been often found
in muddy area or drying lakes of JRI (Bart van de Vijver, personal communication). It cannot
be ruled out, that the lake has possibly dried out completely for some periods with almost no
vegetational growth. Stable conditions are also supported by stability (uniformity) among
diatoms species (higher Evenness ~0.87). Seasonal drying of shallow lakes appears on JRI
also in present time, as has been shown by Váczi et al. (2012).
6.1.3 1520–220 cal. yrs b2k: Filling of the lake, high diatom variability
The next period corresponding with the biozone 3 limits is distinctive by its duration (~1300
years), however based on the abiotic proxies’ outcomes, period was divided into three
subzones (3a–c). Biozone 3a (1520–980 cal. yrs b2k) initiates on high peak (~32) of richness
index at around 1500 cal. yrs b2k, which is followed with its falling course to ~900 cal. yrs
b2k (biozone 3b: 980–460 cal. yrs b2k). Approximately 1100 cal. yrs b2k a small tinge of
peak in the productivity occurs, which later credibly results as one of the factors increasing
TOC content (maximum 1.85 %) in the biozone 3b.
In comparison to biozone 2, subzone 3a contains a slight decrease in Evenness (~0.76),
which imply that diatom assemblages are less balanced, however still diverse. Planothidium
rostrolanceolatum (~0.4) and Sellaphora gracillima (~0.1) forming ~50 % of the diatom
assemblages in the biozone 3a suggest a significant influence of a flowing water in the lake
basin. This observation might have indicated lake water re-filling (or formation of
throughflow lake; Bart van de Vijver, personal communication) associated with high diatom
species variability (high richness, but not necessarily low Evenness). Biozone 3a also begins
with the highest peak of χ, which then starts to decrease (minimum ~500 cal. yrs b2k).
Relatively high χ fits with reduction in sediment input (0.125 mm.yr-1
) lasting throughout all
biozone 3 subzones. However, reduction in sedimentation rate, most probably associated
with frozen water level (colder conditions), is not with a trend arrangement of some other
biotic proxies (e.g. higher richness index in 3a – 26).
52
In comparison with 3a, diatom fauna in 3b is more balanced (Evenness 0.79), however
less diverse (noticeable peak at ~530 cal. yrs b2k) and productive, which unexpectedly does
not correspond with higher TOC content (~1.14 %). TOC variable then could be rather
associated with increased abundance of other organisms (Leloup et al., 2013), such as green
algae and cyanobacteria. The second half of the biozone 3b is also distinctive by increased
presence of Achnanthidium australexigum (~0.1), a species typical for “allegedly warmer”
biozone 1 (Kateřina Kopalová, personal communication), which could mean together with
previous observations slow transition to warmer deflection in the climate, even though there
is still no correspondence with productivity. Decrease of Planothidium rostrolanceolatum,
followed by emergent increase in the biozone 3c (460–220 cal. yrs b2k), is in similarity with
biozone 3a. Most of the remaining proxies (predominantly inorganic chemistry, together
with grain-size distribution), except χFD (continual increase since 2000 cal. yrs b2k) are
found to be oscillating around the same value in all three subzones (a–c), ensuing no
significant course alters.
6.1.4 220 cal. yrs b2k–recent: Recent lake development
Lake development in the last ~220 years is affiliated with abiotic proxy changes, rather than
with diatom assemblage alternations. Increasing χFD, MGS (~9.01 μm), and thus sand
fraction (3.43 %) is probably originating in warmer phase connected with higher meltwater
input and from thawing glaciers and permafrost active layer. This coincides with rapid
advance of sedimentation rate towards the present times. In comparison to previous period,
selected ratio proxies show a slightly higher variability, however no solid conclusion is
possible to obtain due to reduced time resolution for this short time-interval.
6.1.5 Comparison of Lake Anonima proxy data with other palaeoclimatic
reconstructions from wider area
Indication of warmer climatic phase, in Lake Anonima core recorded in the time interval
2440–2050 cal. yrs b2k, is in line with the findings published in several other papers from
JRI region. Mulvaney et al. (2012) calculated temperature (°C; 10 or 100-yr average)
anomalies (relative to 1961–1990 AD average) based on δD measurements, which were
extracted from 363.9 m-long ice core from the top of Mt. Haddington (1630 m), JRI. The ice
53
core extends the longest among the other AP ice records (Thompson et al., 1994; Thomas et
al., 2009), and covers almost 14,000 years of climatic history.
The Holocene temperature history from the JRI ice core is characterized by an earlyHolocene
climatic optimum (from now on a “hypsithermal” term is used) that was 1.3±0.3
°C warmer than present. This early-Holocene temperature shift on the eastern AP is
distinctive by high magnitude and progression, which is in coincidence with trend observed
in ice-core records from the Antarctic continent (Masson-Delmotte et al., 2011). Following
this widespread early-Holocene hypsithermal, temperature trend shows a long interval of
stable climate in the period 9250–2550 yrs b2k (mid-late Holocene hypsithermal). During
this interval, the mean temperature anomaly (0.2±0.2 °C) indicates that conditions at JRI and
around eastern AP were comparable to warm conditions observed at this site over recent
decades. Within this interval of stability, the JRI isotope record indicates that from ~5050 to
3050 yrs b2k conditions may have been only marginally warmer than present. However,
positive temperature anomalies at 10-yrs averages mostly persisted until 2000 yrs b2k and
interchanged with negative temperature anomalies at 10-yrs averages starting from 2380 yrs
b2k (Mulvaney et al., 2012). Therefore, the period 2550–2000 yrs b2k could be regarded as
transition from mid-late Holocene hypsithermal to late Holocene Neoglacial phase.
Termination of the mid-late Holocene hypsithermal in Lake Anonima record is
considered at 2050 cal. yrs b2k, which nearly correspond to 10-yrs average temperature
anomalies from JRI ice-core record (Mulvaney et al., 2012), which significantly positive
temperature anomalies terminates around 2000 yrs b2k. Lake Anonima transition is predated
by lake record from Beak Island (Sterken et al., 2012), situated ~20 km north of Lake
Anonima, where mid-late Holocene hypsithermal is defined within the period 3219–2170
cal. yrs b2k. This phase is by Sterken et al. (2012) characterized by a prolonged peak in
organic carbon content (eq. TOC), total pigment concentrations and an increase in sediment
accumulation rates. According to authors, total diatom concentrations remained nearly
stable, which suggest that this period of higher primary productivity was dominated (in
studied lakes titled BK 1E and BK 1D) by other organisms including green algae, yellowgreen
algae and cyanobacteria representing an increase in the diversity of the phototrophic
community. The reappearance of phaeophorbide a (product of chlorophyll breakdown →
photosensitizer) is likely related with an increased grazing activity by zooplankton and the
development of planktonic community; the latter being supported by the sub-dominance of
54
the tychoplanktonic diatom species Staurosirella pinnata. Higher TOC and diatom
abundance at 2350 cal. yrs b2k in Herbert Sound is also presented by Minzoni et al. (2015).
These observations are in a consistent agreement with biotic data from Lake Anonima core.
On the other side, Sterken et al. (2012) also proved a presence (although in small
numbers) of some typical aerophilic diatoms, which are being in this case probably
associated with the development of moss communities found on the north-western shores of
the lakes, fed by small, braided melt-water streams. The presence of Euglypha genus within
this interval could corroborate this interpretation, pointing to moss growth in the catchment
or littoral area of the lake (Douglas and Smol, 1995; Vincke et al., 2006). Based on Sterken
et al. (2012) and Lake Anonima data presented here, this period is in agreement with a
warmer climate and longer ice-free periods.
The period of increased primary productivity coincides quite consistently (in advance)
with a period of increased diatom abundance in southern Prince Gustav Channel from 4320
to 3580 cal. yrs b2k towards a peak between 2890 and 1390 cal. yrs b2k (Sterken et al.,
2012). In general, it is likely an expression of the mid-late Holocene hypsithermal, a warm
period found in several parts of the AP and elsewhere in Antarctica (Hodgson et al., 2004,
Verleyen et al., 2011). The exact timing of this event is however still inadequately resolved
around the continent (Hodgson et al., 2009), but at nearby JRI it is recorded in lake sediments
between ~4730 and 3170 cal. yrs b2k (4050 and 3040 14
C yrs b2k; Björck et al., 1996),
although the dating resolution of this study is limited to two-three radiocarbon dates per core.
Better-dated records (14 radiocarbon dates) from western AP, shows for the South Shetland
Islands data the mid-late Holocene hypsithermal at ~4430–2580 cal. yrs b2k with an
optimum slightly after 3170 cal. yrs b2k (4050–2580 and 3050 14
C yrs b2k; Björck et al.,
1993). More recent works from this region (Byers and Barton Peninsulas), based on dating
of lake sediments and positions of frontal moraines, however revealed that no significant
glacial advances have occurred during the last 2450 cal. yrs b2k (Oliva et al., 2016), which
might be in correspondence with mid-late Holocene hypsithermal. Chun et al. (2017) proves
that Fildes Peninsula (King George Island, South Shetlands) has experienced mid-late
Holocene hypsithermal between 4450 and 2750 cal. yrs b2k. Similarly, to the north in the
South Orkney Islands well-dated records (13–16 radiocarbon dates) place the mid-late
Holocene hypsithermal between 3850–1450 cal. yrs (Jones et al., 2000; Hodgson and
Convey, 2005). This suggests that the onset of warming at Beak, Vega and James Ross
55
islands might have been buffered by the cooler climate systems of the Weddell Sea Gyre to
the east of the AP. Less precisely dated (due to marine reservoir effects) evidence for a warm
period along the eastern margin of the AP suggests a previous collapse of the Prince Gustav
Channel ending at ~1950 14
C yrs b2k, when the ice shelf reformed (Pudsey and Evans, 2001).
Also, Larsen-A ice shelf fluctuations within its stability/extent were recorded at ~3850, 2150
and 1450 cal. yrs b2k (Brachfeld et al., 2003). It is clear from previous text, that various
proxy evidence exists for a mid-late Holocene hypsithermal on the AP (Bentley et al., 2009),
although there is a significant lack of consensus on its spatiotemporal appearance not only
in JRI region, but also for the whole AP area.
The following period, described by Sterken et al. (2012) as “climate deterioration” with
an onset at 2170 cal. yrs b2k, is consistent with already mentioned significant shifts in most
proxies from the Lake Anonima, however Beak Island records are slightly pre-dating this
episode (circa 120 years) in comparison to our record. This might be affected by the scarcity
of geochronological data around the shift in both studies and the onset of “climate
deterioration” is thus based on the interpolation of underlying and overlying data in both
age-depth models. During this phase, lakes BK1E and BK1D (Beak Island) reveal a rapid
decrease of total pigment concentration and total organic content (in this case loss-onignition
at 550 °C – LOI550, which would correspond well to TOC presented in this study),
which are in addition supplemented by detected presence of Brachysira minor and
Psammothidium abundans. Existence of Brachysira minor on both localities during this
phase enhances concurrent environmental signal consisting of low nutrients, low carbonate
content and reduced productivity. According to Wetzel (2001) this diatom species might
have been additionally associated with the acidic conditions originating in benthic mosses
establishment onto the lake floor. Importance of Brachysira minor is evidenced by its use as
a bioindicator for mentioned specifics also in other parts (lakes) of the AP region (Gibson
and Zale, 2006), however other studies concerning this area set this “Neoglacial” period to
a later start, justifying Weddell Sea Gyre as a main reason of cooling effect. Additionally,
Pudsey and Evans (2001) conclude that around 1950 14
C yrs b2k (2σ: 1730–1994 cal. yrs
b2k) the Prince Gustav Channel has reformed its ice shelf, which is in slightly delayed
correlation with Beak Island and Lake Anonima records. However, due to the usage of the
radiocarbon age instead of calibrated one, comparison caution is on point. This observation
is based on absence of ice-rafted debris (IRD) in the marine sedimentary cores after the midHolocene
(time interval according to cited authors: 5000–2000 14
C yrs b2k). Neoglacial
56
cooling might also be evidenced by ice shelf stability recorded in the Larsen-A region after
1450 14
C yrs b2k (Brachfeld et al., 2003). Oppositely, Minzoni et al. (2015) shows counter
trend consisting of high diatom abundance and TOC content at 1400 cal. yrs b2k.
Exception in “earlier beginning” holds JRI ice core record (Mulvaney et al., 2012), where
the start of this pronounced cooling is registered at 2550 yrs b2k. However, the period 2550–
2000 yrs b2k could be regarded as transitional from late Holocene hypsithermal to
Neoglacial phase. Later on, temperatures between 850 and 450 yrs b2k were on average
0.7±0.3 °C and 1.8 ± 0.3 °C cooler than present on annual and decadal timescale,
respectively. This coincides very fitly with higher values of PC1 (“climate”) score of the
Lake Anonima data (Fig. 17). The support of “earlier Neoglacial onset” is also presented by
Minzoni et al. (2015), where 11-m long marine sedimentary record from Herbert Sound and
Croft Bay, ~27 km to the south-west of Lake Anonima, evidences extended sea cover,
decreased diatom productivity, cold-water diatoms presence (Navicula species, Fursenkoina
curta), altogether with organic carbon content (eq. to TOC presented in this study) and sets
onset of this trend at ~2550 cal. yrs b2k. Important to mention is, that increased IRD is
considered as a result of a local glacier movement (colder period) delivering debris via
iceberg calving on the continental shelf in the ice shelf absent environment. On the other
side, since Pudsey and Evans (2001) operate in Prince Gustav Channel, the area of former
ice shelf presence, they attribute a colder period rather to a decrease of IRD connected with
ice shelf built up.
The previously published research show, that the Neoglacial period has been recognized
in the AP region not as a period of far-reaching glacial advances, like those observed in the
Northern Hemisphere (Badding et al., 2013) and Patagonia (Glasser et al., 2004), but rather
as relatively cooler and drier conditions with suppressed biogenic productivity and increased
sea-ice cover (Leventer et al., 1996; Yoon et al., 2003; Heroy et al., 2008). Spatiotemporal
occurrence of Neoglacial period has been further described also in other studies concerning
adjacent AP regions – e.g., South Shetland Islands (Björck et al., 1991a; 1993), Bransfield
Strait (Fabrés et al., 2000) and East Antarctica (Verleyen et al., 2011).
Further development, corresponding with time interval 1520 cal. yrs b2k – till present,
considered as less variable, reveals most proxies’ changes as rather short-term (decadal,
instead of centennial scale resolution). This period discloses some substantial changes, but
in comparison to rapid shift at 2050 cal. yrs b2k, these changes are, rather short-term with
57
lower signposting ability. Therefore, interpretation and discussion are adapted to and the
observations are thought out with different perspective.
Since 1300 to 430 cal. yrs b2k, Herbert Sound record (Minzoni et al., 2015) shows a longterm
TOC increase together with IRD decrease, suggesting slow, but continual warming. In
addition, subinterval 950–430 cal. yrs b2k in this area (DA1) reveals substantially more
stable and warmer conditions, which might have corresponded with the hint of trend known
as Medieval Climate Anomaly; despite the fact, it is not very prominent in AP region, as
well as in other parts of Antarctica (Tavernier et al., 2014). On the other hand, the Lake
Anonima record shows in the corresponding period (~670–490 cal. yrs b2k) rather colder
conditions evidenced by a PC1 (“climate”) increased values. However, since the Lake
Anonima represents a more-detailed record (in comparison to Minzoni et al., 2015), colder
phase more likely evidences short-term climate fluctuation, sort of which a long-term marine
record cannot capture in a such detail. On the other side, Minzoni et al. (2015) shows an
increase of IRD together with TOC decrease since 430 cal. yrs b2k, but it is not in
chronological correlation with the Lake Anonima record. This is proved by the JRI ice-core
δD record (Mulvaney et al., 2012), which reveals that the strongly negative temperature
anomalies persisted here till 280 yrs b2k, however the coldest period was 650–420 yrs b2k
with continuous temperature increase till present. This partly coincides with the Lake
Anonima record, in which is this warming showed since 500 cal. yrs b2k by negative values
of PC1 “climate” component.
As presented in 6.1.3 relatively high TOC at around 590 cal. yrs b2k does not suggest
warmer conditions, but instead it could be indirectly linked with colder conditions (longer
frozen lake). However, origin of TOC peak might have been explained by concurrent
influence of cyanobacteria and green algae production. Together with decreased
sedimentation rate it implies proportionally higher content of organic matter from planktonic
organisms in seasonally melting lake, which died during the lake re-freezing.
In addition, lake sediments from Beak Island indicate warming with beginning at ~1510
cal. yrs b2k (Sterken et al., 2012), and together these records demonstrate collectively rather
the absence of a wide-spread LIA signal around the AP, which is in general recorded in the
peri-Atlantic sector of the Northern Hemisphere (Mann et al., 2008). Whether LIA in this
region have existed or not, this “colder phase” signal (since 430 cal. yrs b2k) in Herbert
Sound is equally supplemented by high variability in diatom abundance and TOC content
58
since 650 cal. yrs b2k, which might suggest enhanced seasonal variation in sea-ice and ice
rafting.
Climate warming and associated recent ecological changes are marked also in Beak Island
record since 593 cal. yrs b2k by peak of TOC, lower wet density, increase in diatom species
richness and species favouring high nutrient concentrations (e.g. Nitzschia frustulum, N.
inconspicua). From the recent climate perspective, the JRI ice core shows that the recent
phase of warming on the northern AP began in the mid-1920s and that over the past 50 years
the temperature has risen at a rate equivalent to 2.6 ± 1.2 °C per century. However, to address
changes in lake archives for the most recent period, it is required to obtain a high-resolution
data supplemented by firm 210
Pb and 137
Cs (or other type of) chronology.
Due to limited scope as well as time resolution per one sample (10–50 years), and also
with respect to coherence extent of this thesis, discussion concerning linkage of forcing
mechanisms (solar factors, ocean circulation, ENSO etc.) to AP climate change during the
Holocene is not included into consideration. This topic is in more detail discussed by other
authors, e.g. by Bentley et al. (2009). Nevertheless, the sediment core from Lake Anonima,
Vega Island is amongst the first records (Zale and Karlén, 1989; Björck et al., 1991a; 1996;
Ingólfsson et al., 1992; Hjort et al., 1997; Mulvaney et al., 2012; Sterken et al., 2012) to
show a lacustrine response to the most recent period of Holocene environmental variability
within a long-term historical context beyond instrumental records.
6.2 Lake Mathiesondalen 3 evolution in the Late Holocene
6.2.1 700–380 cal. yrs b2k: High sulphur and χ and low TOC
Age-depth model revealed that Lake Mathiesondalen 3 sedimentary record spans the last
~700 years of the environmental history, implying this record as rather short-aged with
possible and thus advantageous low influence of the compaction effect. Not only moredetailed
time resolution, but also age overlap with medieval and modern human times is
considered as beneficial, meaning a possibility to synthesise environmental noesis with
specifics of the human presence in the area. In contrast, short-aged information might
represent different and misleading trend, not fitting into a medium- or long-term scale.
Unchanging conditions, which have occurred during the sedimentation within the lake, lead
to the fact that particular proxy data have enabled to estimate potential main events only.
Notwithstanding data interpretations are also limited to only non-biotic results present here
59
(geochemistry and physical proxies), due to absence of diatom and Cladocera species,
Chironomidae-based reconstruction will otherwise presumably reveal promising results.
Notably in general, samples contain 50–150 chironomids per 1 g of dry sediment (Daniel
Vondrák, personal communication) and result of Chironomidae-based reconstruction will
be presented elsewhere. As a side note, unless stated otherwise, discussed are always both
values (d+w) relating the same proxy.
Evident persistent TOC reduction/stagnation (oscillating around ~5 %) in this period
might signify less progressive conditions, associated rather with slugging organic
productivity (Maher and Thompson, 1999). Proxy for biogenic silica (BSi) estimation Si/Ti
(Brown, 2015) demonstrates as well increasing trend towards the core’s top, with apparent
fluctuation at the bottom, which is in linkage (especially for dried samples) with TOC pattern
suggesting association between organic matter and BSi (Boyle, 2001). Relatively low TC
content (~8.5 %) is in line with comparatively increased χ values implying environment, that
is comparatively low in inorganic carbon (carbonate, CO2, carbonic acids, bicarbonate
anions), but relatively rich on “magnetisable” matter. Although TOC is low, it still holds
major proportion of TC, meaning TIC influence from carbonates within this period is less
significant. Possible explanation of this trend might lie in a rather undeveloped melt-water
streams supplying the lake and non-active permafrost and its active layer, which touts to
increase the TC content in the lake as well (Hendy and Hall, 2006) due to likely low
temperatures (resting in colder period?). Unlike TOC, the TIC steadily increases until the
380 cal. yrs b2k (followed by a rapid drop – 6.2.2) suggesting increasing TIC input from
surrounding lake catchment area. In addition, due to the length of the drilled core, TIC is
unlikeable to be sourced from carbonate bedrock in the lake floor. Instead, carbonates or
evaporates were transported either dissolved, or in small particles by active layer thawing
water from the catchment.
To interpret dominating patterns via PCA results of dried samples (Fig. 16), PC1
positively correlates with organic matter (TOC) as well as with carbonate material (TC,
TIC). On the other side, PC1 shows tight negative relationship with the metal elements and
χ. Therefore, PC1 might represent changes between of autochthonous (positive correlation
– material native to its location) vs. allochthonous (negative correlation – sourced from
catchment area) influence. This assumption would intent rather allochthonous input within
this interval, meaning suppressed processes in a lake body and/or lake’s closest area, and
60
oppositely more developed processes in more distant places influencing the lake.
Allochthonous input is also supported by Al and Si “clustering” within a PCA suggesting a
separate way of transport (origin) in the lake’s composition, which is in agreement with the
fact, that these two elements form (alumino-)silicates, as one of the most common mineral
group in the Earth crust (Cuven et al., 2010). Some claim that Al/Si ratio might reflect
weathering processes (Van Daele et al., 2014), however in the Mathiesondalen’s case this
proportion of dry material reveals rather stable values throughout the profile, indicating static
supply of siliciclastic material to the lake during this time interval. Followingly, support of
allochthonous input could be represented by χ measurements at different frequencies, which
showed that χ15616Hz (high-field) bears at some depths lower values than χ976Hz (low-field).
This implies negative, or extremely low χFD proposing sort of presence of stable single
domain (SSD) magnetic grains instead of superparamagnetic (SP) ones (Dearing et al.,
1996). All in all, these observations are likely in agreement with increased sedimentation
rate in the time interval 700–600 cal. yrs b2k.
However, since an apparent form of bound between TS and Fe (d has tighter correlation)
is appropriately reflected in the PCA’s first components (Fig. 15, 16), interpretation gets
more complicated. To explain, χ reflects Fe content very fitly, which consistently supports
“allochthonous” interpretation from the previous paragraph, furthermore implying increased
denudation and detrital input (Maher and Thompson, 1999). On the other side, as seen from
PC1, the TS content reflects χ and Fe variability as well. This could mean, that TS content
might have separate origin and thus its bound to PC1 is in this case probably very essential,
since sulphur might be part of H2S constituted by reducing bacteria
(decomposition/oxygenation of organic matter; Wilkin and Barnes, 1997). Anoxic/anaerobic
environment conditions (in amictic lakes) may force disulphide to react with Fe to create
iron sulphide (FeS) and iron minerals, such as greigite (Fe3S4) and pyrite (FeS2; Håkanson
and Janson, 1983). However, and even though, magnetic minerals such as greigite are
identified as a source of remanent magnetism in sediments, these deposits may be not only
of inorganic origin, or biologically mediated by sulphate-reducing micro-organisms, but also
of magnetotactic bacteria activity (Mann et al., 1990). Latter concerns mostly sulphid-rich
environment, which implies this factor must be considered as the most likely. Nevertheless,
TS in this period (almost) corresponds with decreasing trend of TOC supporting generally
known correlation of these proxies in oligotrophic lakes (Cohen, 2003). Nutrient-poor
61
amictic and oligotrophic lake type might be supportive in this case when considering diatom
absence causes.
However, sulphur might not be only a part of organic compounds nor H2S, but also of
sulphates and sulphides, such as gypsum (CaSO4 · 2H2O) and anhydrite (CaSO4), which are
known from lake catchment (Dallmann et al., 2004) reflecting allochthonous interpretation.
Another source of sulphur in the sediment within this study might have been associated with
the sea spray process, emitting sulphur into the atmosphere in a form of marine aerosol and
then falling through condensation into the lakes (O’Dowd et al., 1997). Via sea spray, the
sulphur content can generally reach up to 8 % (Leng et al., 2012). However, there are also
several other origins of sulphur in the freshwater lakes in general, such as volcanic eruptions
(Pfeffer et al., 2006) or typhoon activity (Ding et al., 2016). Since these sulphur-enhancing
phenomena have not been described from the studied area yet, they are not considered as
influential.
Unlike in Lake Anonima, Rb/Sr correlation with other abiotic proxies, e.g. TS and χ (PC1)
rather expresses changes in geochemistry (mineralogical or textural changes of mostly
incoming material to the lake; Hošek et al., 2015) instead the climate changes, although even
these conditions cannot be excluded, as they might play less important role. This prevailing
assumption thus gives to the first component also secondary “physical-chemical” attribute,
which induces most of the environment settings within the in-catchment conditions.
Assuming mentioned χ values measured in different frequencies, it cannot be excluded
that presented data may be a result of instrument’s error, or the detection limit. Supportive
fact however is, that χ “wet” and dried samples are differentiated by nearly constant value
(Fig. 12) intending diamagnetic H2O influence as uniform throughout a profile. The H2O
proportion does not decrease towards the lowest section, as it is empirically described
(Tjallingii et al., 2007). Notwithstanding, separate PCAs onto both sample groups (d+w)
proved, that there is nearly no significance of H2O influence onto main dominating patterns.
6.2.2 380 cal. yrs b2k–recent: Variable TOC and TC enhancement
Although it may seem that courses of particular proxies do not reply to environmental
changes, nor these shifts are absent in general due to persistence of stability in the
environment, some parameters’ behaviour could be still attributed to different conditions’
development. Since 380 cal. yrs b2k, the TOC content shows more developed variability
62
possibly supplemented by growing linear trend. TC is reaching its maximum with Ca content
following the same (increasing) pattern (Fig. 11). Oppositely, TS and χ influence slowly
vanishes, meaning probably the establishment of a carbonate-rich environment (possible
CaCO3 precipitation or water flow input of calcareous material; Koinig et al., 2003). Unlike
the previous period (6.2.1), proglacial streams supplying the lake might have been more
dynamic, bringing thus enlarged loads of glacial, i.e. siliciclastic, material (Cecil, 1990).
Apart from 700–380 cal. yrs b2k, this period shows several peaks and drops of certain
proxies, which might be correlated between each other. Specifically, synchronous trend
occurs at around ~380 cal. yrs b2k in Si (d; increase), Ca (d; decrease) and Al (d; decrease)
suggesting potential enhancement of coarser particles or possible BSi increase. Another
synchronic change occurs approximately at ~220 cal. yrs b2k, with TOC (w), Fe/Mn (w),
Si/Ti (w), Zr/Ti (w) decrease and Al/Si (w), Si (d), Al (d+w), Ti (d) increase. Very similar
shifts occur also at around ~50 cal. yrs b2k. Cause of such a change may arise from sudden
material input, or also from in-lake processes. However, these assumptions need to be
supported by better understanding of the environment based on further, especially biological
data.
Figure 15 Lake Mathiesondalen 3: PCA 1 (two-dimensional) biplot plotting PC1 (40.40 %) against PC2 (17.81
%) along with all abiotic variables being extracted from “wet” (w) samples and colourised by data type.
63
Figure 16 Lake Mathiesondalen 3: PCA 2 (two-dimensional) biplot plotting PC1 (53.07 %) against PC2 (12.20
%) along with all abiotic variables being extracted from dried (d) samples and colourised by data type.
6.2.3 Comparison of Lake Mathiesondalen 3 proxy data with other palaeoclimatic
reconstructions from wider area
Palaeoenvironmental evolution in the area has been, from perspective of oceanic circulation
patterns, influenced by enhanced water inflow from the Atlantic Ocean since ~1000 yrs b2k,
which have later resulted in generally warming conditions, though the trend conditions differ
in detail. This course was marked in numerous marine (Rasmussen et al., 2014; Werner et
al., 2016) and lacustrine (D’Andrea et al., 2012; Balascio et al., 2018) records.
Specifically, D’Andrea et al. (2012), as one of the most suitable research to be compared
with, show sub-decadal- to multidecadal-scale record of summer (June, July, August – JJA)
temperatures for the past 1800 yrs b2k from lake sediments of Kongressvatnet on the West
Spitsbergen. This study reveals, that JJA temperatures since ~700 yrs b2k till present were
in this area more variable, but still with ongoing warming trend. Before this period, an
interval ~1100–700 yrs b2k has confirmed rather colder conditions (supported by Humlum
et al., 2005 in central Svalbard), which could have been sort of precursor to the possible LIA.
Time (sub-)interval within the range ~750–700 yrs b2k has been typical by significantly
colder summer temperatures, which corresponds to extensive alpine glaciation on Svalbard
in some areas, already described by Baranowski and Karlén (1976) and Werner (1993). This
64
implies that reduced ablation (due to generally lower summer temperatures) contributed to
glacier advances at this time, while oppositely following increased summer temperatures
have become a limitation of these advances.
Authors claim that increased winter precipitation contributed to glacier advance more
than generally low temperatures (Svendsen and Mangerud, 1997; Humlum et al., 2005) at
the turn of the 18th
and 19th
century. This explanation is associated with North Atlantic
Oscillation changes coming out from correlation between winter precipitation and North
Atlantic Oscillation; Dickson et al., 2000). Nevertheless, overall warming trend has been
recognized and attributed to the amount of Atlantic water transported towards the north by
CDW, which in other words means that proximity of WSC plays an essential role despite a
probable delay of warming distribution in west-east direction.
Another record from north-western part of Svalbard is analysed by Røthe et al. (2015)
and its observations are in contrary to the D’Andrea et al. (2012). This research is focused
on local glacier’s equilibrium line altitude (ELA) reconstruction and reveals, that local
climate conditions were in favour to a glacier advance at the time around ~1200 cal. yrs b2k.
In the subsequent period 1200–900 cal. yrs b2k, ELA has oppositely risen, which resulted in
glacier retreat, however unlike in D’Andrea et al. (2012) and according to other authors, LIA
period lasted here since 500 cal. yrs b2k till the beginning of the 20th
century (lowest position
of ELA took place at: 450, 280, 180 cal. yrs b2k). On the other side warmer conditions,
which were oppositely reported in interval 275–185 cal. yrs b2k (Røthe et al., 2015), are in
correlation with enhanced warming during mid-18th
century described by D’Andrea et al.
(2012). Followingly, Røthe et al. (2015) claim the position of local glacier’s ELA steadily
rose since 100 cal. yrs b2k, which is in line with 20th
century warming trend. Another record
by van der Bilt et al. (2015) from north-western Svalbard supports glacier presence/advance
(LIA decrease) after 750 cal. yrs b2k, which is in agreement with Røthe et al. (2015).
Despite the fact, that LIA has been described widely in the Arctic (Solomina et al., 2015),
Lake Mathiesondalen 3 record does not seem to respond to such a phenomenon in a welldefined
and full extent during the period after 700 cal. yrs b2k as well as at the turn of the
18th
and 19th
centuries. On the other side, proxy dataset does not seem to show an absence
of these cold conditions, being probably expressed in this context as LIA. Specifically,
interval 700–380 cal. yrs b2k shows TOC presence as comparatively low, which might have
suggested mild phase within colder conditions associated with glacier advances due to
65
increased precipitation during a winter season, already described by D’Andrea et al. (2012)
in the earlier interval. Assuming these colder conditions, allochthonous input might be
supportive, when suggesting rather increased input from the upper part of the catchment
triggered by glacier movement. Associated winter precipitations (inducing glacier advance)
rather than reduced summer ablation (due to lower summer temperatures) might be positive
when considering increase of magnetisable matter and TS. This could have been linked with
lower summer temperatures creating frozen water lake level and thus rather oligotrophic
(anaerobic) conditions in the lake supporting formation of sulphur compounds as described
before (6.2.1). In addition, and also, increased precipitation inflow does not have to
necessarily contradict with claimed “colder phase”, which might have persisted until 380
cal. yrs b2k.
Increasing TOC, together with Ca increased share in the interval 380 cal. yrs b2k till
recent can be correlated with higher productivity associated with ongoing warming and
moistening, which is in line with higher influence of autochthonous productivity in the lake
and surrounding catchment. As in the previous phase, these warmer conditions might
correlate, although with certain time-lag with observed variable warming trend since 700
cal. yrs b2k presented by D’Andrea et al., 2012). Followingly, when considering dominant
factor influencing climate-glacier behaviour, presented observations indicate, that WSC
might play an important role as it brings warm and moist air from lower latitudes. This also
suggests, that position of Lake Mathiesondalen 3 in central Svalbard could explain discussed
and rather delayed onset of warming. This trend, although less clear and with earlier onset
(~800 cal. yrs b2k), is supported also by diatom assemblages (Amphora affinis, Citellus
richardsonii) in ~7 km distant (NW) Lake Garmaksla analysed by Roman (2014, 2017).
The record from Lake Garmaksla was supposed to be essential for synthesis/comparison,
however after a stratigraphic correlation of three differences based on sequence slotting
(Thompson et al., 2012) the lowermost section revealed to be distinctive by circa 5800 years
(cal. yrs b2k), enabling thus potentially the comparison only with the topmost (mesotrophic)
section. Nevertheless, biostratigraphy and cluster analyses of diatom communities defined
five biozones (A–E) in total (Pinseel, 2014), with the upmost D (~870–450 cal. yrs b2k) and
E (~450 cal. yrs b2k – present) zones nearly entirely overlapping Mathiesondalen’s record.
Assuming the correct age-depth modelling and presence of diatom biota in general, it is not
excluded, that Mathiesondalen’s biostratigraphy analysis would not have possibly reveal
66
similar results. After all, and as discussed previously, abiotic proxies slightly suggest minor
(but substantial?) changes within the last ~380 years, which could have been associated with
D to E biozones transition at around 450 cal. yrs b2k in Lake Garmaksla as well. However,
these observations and assumptions need to be supplemented by further biostratigraphic
data.
Presented research and observations suggest, that majority of available records present
rather a widespread glacier growth evidence during the LIA, or glacier advances at the turn
of the 18th
and 19th
century (Rachlewicz et al., 2007), although some studies have suggested
just the opposite (D’Andrea et al., 2012; Spielhagen et al. 2011). Nevertheless, both
Mathiesondalen 3 and Garmaksla records are shown to be very unique when studying a lake
response to these widespread, but specific phenomena. This however implies that the
climatic driver of glacier growth on Svalbard remains still elusive and thus deserves future
investigation. We can only speculate about the effect of increasing precipitations, which
would lead to positive mass balance of local glaciers even during phases of slow atmospheric
warming. This might be the case of the turn of the 18th
and 19th
century. However, the
continuous warming trend lead to subsequent glacier shrinkages, which started in the area
from the end of the 19th
century (Rachlewicz et al., 2007) and accelerates towards present
(Małecki, 2016).
6.3 “Comparing the evolution of North and South since 700 cal. yrs b2k”
Previously presented outcomes based on a two-core palaeolimnological research have been
used to set up a basis for a discussion, considered as the “state-of-the-art” component of this
thesis. In comparison to the Lake Mathiesondalen 3 record, the Lake Anonima record shows
more pronounced long-term changes, which could be attributed to a differently-parametrized
and dominantly prevailing factors influencing the record. Also, as seen from proxy data
synthesis (Fig. 17) the Lake Anonima record’ resolution is less-detailed, implying more
complicated capture of the short-term changes, which leads to a conclusion, that it is far
more difficult to obtain solid and correct palaeoreconstruction. Followingly, due to the Lake
Mathiesondalen 3 age, only the last 700 cal. yrs b2k are possible to be compared and also,
due to a limited scope of this thesis, together with inability to compare all proxies in their
best complexity and detail, author decided to describe only selected (most influential)
variables, which explain behaviour throughout the major part of the overlapping time period.
67
Specifically, sedimentation rate, main PC’s and TOC were chosen for comparison due to
their most assuring ability in the context of comparison of environmental and climatic
changes.
Sedimentation rate together with PCA’s results were preferred in order to obtain complex
image about the dynamics influencing the lake’s catchment. To compare, average
sedimentation rate for the last 700 cal. yrs b2k varies between cores within several orders.
In Lake Anonima, a mean sedimentation rate is ~0.42 mm.yr-1
, while in the Lake
Mathiesondalen 3 this parameter reaches up to 5.66 mm.yr-1
, meaning much faster material
transport in the selected Arctic catchment. However, influence of in situ material together
with compaction effect needs to be considered as well. Even though the sedimentation rate
highly depends on many factors (i.e., glacier and hydrometeorological activity), the latter
does not seem to correspond with generally known (Wolfe et al., 2004) extremely low
sedimentation rates in Polar areas. Another variable, which reveals most of the core’s
behaviour is the “climate” component (explaining 61.73 % of variability) in the Lake
Anonima and the first “allochthonous vs. autochthonous” component (explaining 53.07 %
of variability) in the Lake Mathiesondalen 3. Both components’ behaviour together with
previously discussed sedimentation rate character explain main changes during the
sedimentation since 700 cal. yrs b2k. This is properly supplemented by TOC proxy, which
is considered as main indicator of primary productivity in the lake. Besides, the differences
in principal component analysis explanation show that the Lake Anonima record is
climatically affected and might bring regional interpretation of climatic and environmental
changes. On the contrary, the Lake Mathiesondalen 3 record is affected principally by incatchment
and in-lake processes and explains predominantly changes of local environmental
conditions.
Chronologically, development in the Lake Anonima is in this comparison initiated by
small hint of “climate component” positive peak between ~670–490 cal. yrs b2k, suggesting
indication of relatively colder period in the Lake Anonima area. This non-local (catchment),
but rather regional climate pattern is also in agreement with low temperature (°C) anomalies
vacillating between 0.20 and –1.79 °C (mean –0.71 °C), which were extracted from JRI
deuterium-excess ice core record (Mulvaney et al., 2012). Subsequently, sedimentation rate
in this interval is rather low, starting on ~0.1 mm.yr-1
at 700 cal. yrs b2k and slowly rising
up to 0.205 mm.yr-1
at around 580 cal. yrs b2k. Even though low sedimentation rates are in
68
line with the previous observations, better time resolution of material input chronology
would elucidate, whether the recorded “colder” peak is, or is not substantial and marked also
by other proxies. In contrast, increased organic productivity expressed by TOC does not
seem to correlate with the supposed colder peak. As described in 6.1.3 high TOC might not
strictly link towards high (diatom) productivity, but rather to other organisms, such as green
algae and/or cyanobacteria (Leloup et al., 2013) as well as to frozen lake level and
proportionally less accumulated material due to lower sedimentation rate. On the contrary,
the Lake Mathiesondalen 3 reveals, that the whole period since 700 till 380 cal. yrs b2k is
typified by quite unique conditions. As presented before (6.2.3), proxies configuration might
be probably associated with LIA lag, due to the delayed influence of WSC on Central
Spitsbergen climate. Precipitation increase was essential, as it had dominant influence on
glacier growth and input of allochthonous input into the lake, while generally lower
(summer) temperatures were linked with TOC stagnation/decrease and TS together with χ
enrichment via developed anoxic conditions. Increased sedimentation rate (~25 and
decreasing to ~2 mm.yr-1
) at around 700 cal. b2k might be supportive when considering
increased allochthonous (siliclastic) input during this period.
Subsequently, development in the Lake Anonima between ~490 and 340 cal. yrs b2k as
well as since ~300 cal. yrs b2k till present, shows rather the negative scores of the “climate
component”, implying climate warming effect on the sedimentary record. The warming
trend might have been interrupted by a short colder phase between ~340 and 300 cal. yrs
b2k, as seen from Fig. 17. TOC parameter continuously decreases since ~590 cal. yrs b2k,
and it does not probably respond to “climate component” variations since this time, neither.
Also, sedimentation rate persists around 0.2 mm.yr-1
, meaning not well developed response
to the “climate component”. Nevertheless, during the last ~30 cal. yrs b2k this parameter
steeply rises up to nearly 1 mm.yr-1
. The synchronous period (since 380 cal. yrs b2k) is in
the Lake Mathiesondalen 3 marked by prevailing in-lake productivity, which might be
properly supplemented by higher (but variable?) TOC content, as well as the TIC content
affecting the general increase of total carbon towards the present. Sedimentation rate also
rises up to 10 mm.yr-1
. This trend might reflect local climate amelioration.
To meet the aims of this chapter, and also to sum up and highlight the main patterns
describing the proxy variability within studied cores, it is evident, that during the (at least)
last 300 cal. yrs these particular localities have shown similar environmental (warming)
69
conditions. Additionally, and in comparison, to the Lake Anonima (Vega Island), the Lake
Mathiesondalen 3 seems to reveal more dynamic and more rapidly changing environment,
although this sentence needs to be taken into consideration with certain standoff, since
enormous number of variables plays an essential role within a palaeoenvironmental
“recording” in the lake sediments. However, generally it can be stated that the Lake Anonima
record reflects principally the climate variability of Maritime Antarctica over the last two
and half millennia. On the contrary, the Lake Mathiesondalen 3 record shows mainly
variable environmental conditions within the lake and its catchment.
70
Figure 17 Proxy data synthesis (left to right): Mt. Haddington (JRI) Ice-core anomaly (°C) 10 yr. avg. (1961–1990; Mulvaney et al., 2012); Lake Anonima: PC1 (climate);
MGS (µm); χFD (%) and TOC (%). Lake Mathiesondalen 3: PC1 (allochthonous input vs. autochthonous production); TC (%); TOC (%); TS (%) and χ976Hz.
71
7 SUMMARY AND CONCLUSIONS
In this thesis, author has focused on a palaeolimnological research in the Polar areas, while
selected regions (Vega Island, Antarctica; Mathiesondalen, Svalbard), have been subjected
to a more-detailed view. In the literature review chapter, a comprehensive summary based
on a physical-geographical description and published researches from these areas, was
appropriately delivered. Presented palaeoenvironmental studies, which used multiple
techniques and records (ice cores, marine and lacustrine sediments, geomorphology) for
reconstruction, were mentioned and properly introduced.
Main part of the thesis was aimed to give information about the palaeolimnological
evolution in two areas of Polar regions – Lake Anonima, Vega Island in Antarctica and Lake
Mathiesondalen 3, Spitsbergen in Svalbard (Arctic). Using a multi-proxy approach
(petrophysical and in/organic geochemistry) and robust chronostratigraphic framework
based on radiocarbon (14
C) and short-lived isotope (210
Pb, 137
Cs) absolute dating, which was
supplemented by adequate statistical methods (PCA), author was able to extract information
about the dominating patterns driving the environmental changes in selected polar areas.
The presented palaeoenvironmental record of Lake Anonima covers the last 2400 cal. yrs
b2k and captures the termination of the Late Holocene Hypsithermal, together with its
following shifts towards Neoglacial period and recent warming. These phases have been
evidenced by diatom community changes as well as other abiotic proxies. Younger
sedimentary record from the Lake Mathiesondalen 3 includes the last 700 cal. yrs b2k and
based on abiotic proxies it provides unique decadal archive of the local catchment evolution
in the Mathiesondalen, Central Spitsbergen.
In the final chapter, the main differences based on the most valuable proxy data (PC1,
TOC, sedimentation rate), were established. Specifically, it has been concluded that the Lake
Anonima record reflects principally the climate variability of Maritime Antarctica over the
last two and half millennia, but on the other hand, the Lake Mathiesondalen 3 record shows
mainly variable environmental conditions within the lake and its catchment.
72
REFERENCES
Monographies, papers from scientific journals and theses
Aagaard, K., Foldvik, A., Hillman, S. R. (1987): The West Spitsbergen Current: Disposition
and water mass transformation. Journal of Geophysical Research, 92, č. C4, s. 3778–
3784.
Ambrožová, K., Láska, K. (2016): Změny teploty vzduchu na ostrově Jamese Rosse v
kontextu Antarktického poloostrova. In: Nováček, A. [eds.]: Sborník příspěvků Výroční
konference České geografické společnosti. Jihočeská univerzita v Českých Budějovicích,
České Budějovice, s. 20–25.
Allen, J. R. L., Thornley, D. M. (2004): Laser granulometry of Holocene estuarine silts:
effects of hydrogen peroxide treatment. The Holocene, 14, č. 2, s. 290–295.
Appleby, P. G., Oldfield, F., Thompson, R., Huttenen, P., Tolonen, K. (1979): 210
Pb dating
of annually laminated lake sediments from Finland. Nature, 280, č. 5717, s. 53–55.
Aristarain, A. J., Jouzel, J. Pourchet, M. (1986): Past Antarctic Peninsula climate (1850–
1980) deduced from an ice core isotope record. Climatic Change, 8, č. 1, s. 69–89.
Badding, M. E., Briner, J. P., Kaufman, D. S. (2013): 10
Be ages of late Pleistocene
deglaciation and Neoglaciation in the north-central Brooks Range, Arctic Alaska. Journal
of Quaternary Science, 28, č. 1, s. 95–102.
Balascio, N. L., D’Andrea, W. J., Gjerde, M., Bakke, J. (2018): Hydroclimate variability of
High Arctic Svalbard during the Holocene inferred from hydrogen isotopes of leaf waxes.
Quaternary Science Reviews, 183, č. 5, s. 177–187.
Baranowski, S., Karlén, W. (1976): Remnants of Viking Age tundra in Spitsbergen and
northern Scandinavia. Geografiska Annaler, 58, č. 1–2, s. 35–39.
Bennett, K. D. (1996): Determination of the number of zones in a biostratigraphical
sequence. New Phytologist, 132, č. 1, s. 155–170.
Bentley, M. J., Fogwill, C. J., Kubik, P. W., Sugden, D. E. (2006): Geomorphological
evidence and cosmogenic Be/Al exposure ages for the Last Glacial Maximum and
deglaciation of the Antarctic Peninsula Ice Sheet. Geological Society of America
Bulletin, 118, č. 9–10, 1149–1159.
Bentley, M. J., Hodgson, D. A., Smith, J. A., Cox, N. J. (2005a): Relative sea level curves
for the South Shetland Islands and Marguerite Bay, Antarctic Peninsula. Quaternary
Science Reviews, 24, č. 10–11, s. 1203–1216.
Bentley, M. J., Hodgson, D.A., Smith, J.A., Ó Cofaigh, C., Domack, E.W., Larter, R.D.,
Roberts, S.J., Brachfeld, S., Leventer, A., Hjort, C., Hillenbrand, C-D., Evans, J. (2009):
Mechanisms of Holocene palaeoenvironmental change in the Antarctic Peninsula region.
The Holocene, 19, č. 1, s. 51–69.
73
Björck, S., Håkansson, H, Olssen, S., Banekow, L., Janssens, J. (1993): Paleoclimatic studies
in south Shetland islands, Antarctica, based on numerous stratigraphic variables in lake
sediments. Journal of Paleolimnology, 8, č. 3, s. 233–272.
Björck, S., Håkansson, H., Zale, R., Karlén, W., Jönssen, B. L. (1991a): A late Holocene
lake sediment sequence from Livingston Island, south Shetland islands, with
palaeoclimatic implications. Antarctic Science, 3, č. 1, s. 61–72.
Björck, S., Olsson, S., Ellis-Evans, C., Håkansson, H., Humlum, O., de Lirio, J.M. (1996):
Late Holocene palaeoclimatic records from lake sediments on James Ross island,
Antarctica. Palaeogeography, Palaeoclimatology, Palaeoecology, 121, č. 3–4, s. 195–220.
Bjornsson, H., Gjessing, Y., Hamran, S. E., Hagen, J. O., Liestol, O., Palsson, F., Erlingsson,
B. (1996): The Thermal regime of sub-polar glaciers mapped by multi-frequency radioecho
sounding. Journal of Glaciology, 42, č. 140, s. 23–32.
Blauww, M. (2010): Methods and code for ‘classical’ age-depth modeling of radiocarbon
sequences. Quaternary Geochronology, 5, č. 5, s. 512–518.
Blaszczyk, M., Jania, J. A., Hagen, J. O. (2009): Tidewater glaciers of Svalbard: recent
changes and estimates of calving fluxes. Polish Polar Research, 30, č. 2, 85–142.
Blott, S. J., Pye, K. (2001): GRADISTAT: a grain size distribution and statistics package for
the analysis of unconsolidated sediments. Earth Surface Processes and Landforms, 26, č.
11, s. 1237–1248.
Böning, P., Bard, E., Rose, J. (2007): Toward direct, micron-scale XRF elemental maps and
quantitative profiles of wet marine sediments. Geochemistry, Geophysics, Geosystems,
8, č. 5, s. 1–14.
Boyle, J. F. (2001): Inorganic geochemical methods in paleolimnology. In: Last, W. M.,
Smol, J. P. [eds.]: Tracking Environmental Change Using Lake Sediments. Volume 2:
Physical and Geochemical Methods. Developments in Paleoenvironmental Research,
Kluwer, Dordrecht, s. 83–142.
Bradley, R. S. (2015): Paleoclimatology: Reconstructing Climates of the Quaternary. 3rd
edition. Academic Press, Amsterdam, 675 s.
Brachfeld, S., Domack, E., Kissel, C., Laj, C., Leventer, A., Ishman, S., Gilbert, R.,
Camerlenghi, A., Eglinton, L. B. (2003): Holocene history of the Larsen-A ice shelf
constrained by geomagnetic paleointensity dating. Geology, 31, č. 9, s. 749–752.
Brown, E. T. (2015): Estimation of Biogenic Silica Concentrations Using Scanning XRF:
Insights from Studies of Lake Malawi Sediments. In: Croudace, I. W., Rothwell, R. G.
[eds.]: Micro-XRF Studies of Sediment Cores: Applications of a Non-Destructive Tool
for the Environmental Sciences, Developments in Paleoenvironmental Research.
Springer, Dordrecht, s 267–278.
74
Cecil, C. B. (1990): Paleoclimate controls on stratigraphic repetition of chemical and
siliciclastic rocks. Geology, 18, č. 6, s. 533–536.
Chaparro, M. A. E., Chaparro, M. A. E., Córdoba, F. E., LeComte, K. L., Gargiulo, J. D.,
Barrios, A. M., Urán, G. M., Manograsso Czalbowski, N. T., Lavat, A., Böhnel, H. N.
(2017): Sedimentary analysis and magnetic properties of Lake Anonima, Vega Island.
Antarctic Science, 29, č. 5, s. 429–444.
Chen, J., Zhisheng, A., Head, J. (1999): Variation of Rb/Sr ratios in the loess-paleosol
sequences of central China during the last 130,000 years and their implications for
monsoon paleoclimatology. Quaternary Research, 51, č. 3, s. 215–219.
Chun, Z., Sun, L., Wang, Y., Huang, T., Zhou, X. (2017): Depositional environment and
climate changes during the Holocene in Grande Valley, Fildes Peninsula, King George
Island, Antarctica. Antarctic Science, 29, č. 6, s. 545–554.
Clapperton, C. M., Sugden, D. E. (1982): Late Quaternary glacial history of George VI
Sound area, West Antarctica. Quaternary Research, 18, č. 3, s. 243–267.
Cohen, A. S. (2003): Paleolimnology: the history and evolution of lake systems. Oxford
University Press, New York, 500 s.
Cook, A. J., Vaughan, D. G. (2010): Overview of areal changes of the ice shelves on the
Antarctic Peninsula over the past 50 years. The Cryosphere, 4, č. 1, s. 77–98.
Cuven, S., Francus, P., Lamoureux, S. F. (2010): Estimation of grain size variability with
micro X-ray fluorescence in laminated lacustrine sediments, Cape Bounty, Canadian
High Arctic. Journal of Paleolimnology, 44, č. 3, s. 803–817.
D’Andrea, W. J., Vaillencourt, D. A., Balascio, N. L., Werner, A., Roof, S. R., Retelle, M.,
Bradley, R. S. (2012): Mild Little Ice Age and unprecedented recent warmth in an 1800
year lake sediment record from Svalbard. Geology, 40, č. 11, s. 1007–1010.
De Wet, G. A., Balascio, N. L., D’Andrea, W. J., Bakke, J., Bradley, R. S., Perren, B. (2018):
Holocene glacier activity reconstructed from proglacial lake Gjøavatnet on
Amsterdamøya, NW Svalbard. Quaternary Science Rewievs, 183, č. 5, s. 188–203.
Dearing, J. A., Dann, R. J. L., Hay, K., Lees, J. A., Loveland, P. J., Maher, B. A., O’Grady,
K. (1996): Frequency-dependent susceptibility measurements of environmental
materials. Geophysical Journal International, 124, č. 1, s. 228–240.
Dickson, R. R., Osborn, T. J., Hurrell, J. W., Meincke, J., Blindheim, J., Ådlandsvik, B.,
Vinje, T., Alekseev, G., Maslowski, W. (2000): The Arctic Ocean response to the North
Atlantic Oscillation: Journal of Climate, 13, č. 15, s. 2671–2696.
Ding, X., Li, D., Zheng, L., Bao, H., Chen, H. F., Kao, S. J. (2016): Sulfur Geochemistry of
a Lacustrine Record from Taiwan Reveals Enhanced Marine Aerosol Input during the
Early Holocene. Scientific Reports, 6, č. 38989, s. 1–9.
75
Domack, E. W., Foss, D. J. P., Syvitski, J. P. M., McMclennen, C. E. (1994): Transport of
suspended particulate matter in an Antarctic Fjord. Marine Geology, 121, č. 3–4, s. 161–
170.
Domack, E. W., Leventer, A., Root, S., Ring, J., Williams, E., Carlson, D., Hirshorn, E.,
Wright, W., Gilbert, R., Burr, G. (2003b): Marine sedimentary record of natural
environmental variability and recent warming in the Antarctic Peninsula. In: Domack, E.
W., Leventer, A., Burnett, A., Bindschadler, R., Convey, P., Kirby, M. [eds.]: Antarctic
Peninsula Climate Variability: Historical and Paleoenvironmental Perspectives.
American Geophysical Union, Washington, D.C., s. 205–222.
Douglas, M. S. V., Smol, J. P. (1995): Paleolimnological significance of observed
distribution patterns of chrysophyte cysts in arctic pond environments. Journal of
Paleolimnology, 13, č. 1, 1–5.
Dylmer, C. V., Giraudeau, J., Eynaud, F., Husum, K., De Vernal, A. (2013): Northward
advection of Atlantic water in the eastern Nordic Seas over the last 3000 yr. Climate of
the Past, 9, č. 4, s. 1505–1518.
Engel, Z., Nývlt, D., Láska, K. (2012): Ice thickness, areal and volumetric changes of Davies
Dome and Whisky Glacier in 1979–2006 (James Ross Island, Antarctic Peninsula).
Journal of Glaciology, 58, č. 211, s. 904–914.
Ermolin, E., De Angelis, H., Skvarca, P. (2002): Mapping of permafrost on Vega Island,
Antarctic Peninsula, using aerial photography and satellite image. Annals of Glaciology,
34, č. 1, s. 184–188.
Evans, M., Heller, F. (2003): Environmental Magnetism: Principles and applications of
enviromagnetics. Vol. 86. Academic press, New York, 417 s.
Evans, J. A., Pudsey, C. J., Ó Cofaigh, C., Morris, P., Domack, E. (2005): Late Quaternary
glacial history, flow dynamics and sedimentation along the eastern margin of the
Antarctic Peninsula Ice Sheet. Quaternary Science Reviews, 24, č. 5–6, s. 741–774.
Fabrés, J., Calafat, A. M., Canals, M., Bárcena, M. A., Flores, J. A. (2000): Bransfield basin
fine grained sediments: Late Holocene sedimentary processes and oceanographic and
climatic conditions. The Holocene, 10, č. 9, s. 703–718.
Feyling-Hanssen, R. W. (1955): Stratigraphy of the marine Late-Pleistocene of Billefjorden,
Vestspitsbergen. Norsk Polarinstitutt, Oslo, 186 s.
Folk, R. L., Ward, W. C. (1957): Brazos River bar [Texas]; a study in the significance of
grain size parameters. Journal of Sedimentary Research, 27, č. 1, s. 3–26.
Førland, E. J., Benestad, R., Hanssen-Bauer, I., Haugen, J. E., Skaugen, T. E. (2011):
Temperature and precipitation development at Svalbard 1900–2100. Advances in
Meteorology, 2011, č. 1, s. 15.
76
Gersonde, R., Abelmann, A., Brathaeur, U., Becquey, S., Bianchi, C., Cortese, G., Grobe,
H., Kuhn, G., Niebler, H. S., Segl, M., Sieger, R., Zielinski, U., Futterer, D. K. (2003):
Last glacial sea surface temperatures and sea-ice extent in the Southern Ocean (AtlanticIndian
sector): a multiproxy approach. Paleoceanography and Paleoclimatology, 18, č. 3,
s. 1–18.
Gibas, J., Rachlewicz, G., Szczuciński, W. (2005): Application of DC resistivity soundings
and geomorphological surveys in studies of modern Arctic glacier marginal zones,
Petuniabukta, Spitsbergen. Polish Polar Research, 26, č. 4, s. 239–258.
Gibson, J. A. E., Zale, R. (2006): Holocene development of the fauna of Lake Boeckella,
northern Antarctic Peninsula. The Holocene, 16, č. 5, s. 625–634.
Gjerde, M., Bakke, J., D’Andrea, W. J., Balascio, N. L., Bradley, R. S., Vasskog, K.,
Ólafsdottir, S., Røthe, T. O., Perren, B. B., Hormes, A. (2018): Holocene multi-proxy
environmental reconstruction from lake Hakluytvatnet, Amsterdamøya Island, Svalbard
(79.5°N). Quaternary Science Reviews, 183, č. 5, s. 164–176.
Glasser, N. F., Harrison, S., Winchester, V., Aniya, M. (2004): Late Pleistocene and
Holocene palaeoclimate and glacier fluctuations in Patagonia. Global and Planetary
Change, 43, č. 1–2, s. 79–101.
Hagen, J. O., Kohler, J., Melvold, K., Winther, J. G. (2003a): Glaciers in Svalbard: mass
balance, runoff and freshwater flux. Polar Research, 22, č. 2, s. 145–159.
Håkanson, L., Jansson, M. (1983): Principles of Lake Sedimentology. Springer, Berlin, 316
s.
Hamilton, G. S., Dowdeswell, J. A. (1996): Controls on glacier surging in Svalbard. Journal
of Glaciology, 42, č. 140, s. 157–168.
Harland, W. B. (1997): The Geology of Svalbard. The Geological Society, London, 539 s.
Haus, N. W., Wilhelm, K. R., Bockheim, J. G., Fournelle, J., Miller, M. (2016): A case for
chemical weathering in soils of Hurd Peninsula, Livingston Island, South Shetland
Islands, Antarctica. Geoderma, 263, č. 3, s. 185–194.
Hendey, N. I. (1964): An introductory account of the smaller algae of British coastal waters.
Fishery Investigation Series. Part 5: Bacillariophyceae (Diatoms). Her Majesty’s
Stationery Office, London, 317 s.
Hendy, C. H., Hall, B. L. (2006): The radiocarbon reservoir effect in proglacial lakes:
Examples from Antarctica. Earth and Planetary Science Letters, 241, č. 3–4, s. 413–421.
Hennekam, R., Lange, G. (2012): X-ray fluorescence core scanning of wet marine
sediments: methods to improve quality and reproducibility of high-resolution
paleoenvironmental records. Limnology and Oceanography: Methods, 10, č. 12, s. 991–
1003.
77
Hernández-Molina, F. J., Larter, R. D., Rebesco, M., Maldonado, A. (2006): Miocene
reversal of bottom water flow along the Pacific Margin of the Antarctic Peninsula:
stratigraphic evidence from a contourite sedimentary tail. Marine Geology, 228, č. 1–4,
s. 93–116.
Heroy, D. C., Anderson, J. B. (2005): Ice-sheet extent of the Antarctic Peninsula region
during the Last Glacial Maximum (LGM) – insights from glacial geomorphology.
Geological Society of America Bulletin, 117, č. 11–12, s. 1492–1512.
Heroy, D. C., Sjunneskog, C., Anderson, J. B. (2008): Holocene climate change in the
Bransfield Basin, Antarctic Peninsula: evidence from sediment and diatom analysis.
Antarctic Science, 20, č. 1, s. 69–87.
Hjort, Ch., Ingólfsson, O., Möller, P., Lirio, J. M. (1997): Holocene glacial history and sealevel
changes on James Ross Island, Antarctic Peninsula. Journal of Quaternary Science,
12, č. 4, 259–273.
Hodgson, D.A., Abram, N., Anderson, J., Bargelloni, L., Barrett, P., Bentley, M.J., Bertler,
N.A.N., Chown, S., Clarke, A., Convey, P., Crame, A., Crosta, X., Curran, M., di Prisco,
G., Francis, J.E., Goodwin, I., Gutt, J., Massé, G., Masson-Delmotte, V., Mayewski, P.A.,
Mulvaney, R., Peck, L., Pörtner, H.-O., Röthlisberger, R., Stevens, M.I., Summerhayes,
C.P., van Ommen, T., Verde, C., Verleyen, E., Vyverman, W., Wiencke, C., Zane, L.
(2009): Antarctic climate and environment history in the pre-instrumental period. In:
Turner, J., Convey, P., di Prisco, G., Mayewski, P.A., Hodgson, D.A., Fahrbach, E.,
Bindschadler, R., Gutt, J. [eds.]: Antarctic Climate Change and the Environment.
Scientific Committee for Antarctic Research, Cambridge, s. 115–182.
Hodgson, D. A., Convey, P. (2005): A 7000-Year record of Oribatid Mite communities on
a Maritime-Antarctic island: responses to climate change. Arctic Antarctic and Alpine
Research, 37, č. 2, s. 239–245.
Hodgson, D. A., Doran, P., Roberts, D., McMinn, A. (2004): Paleolimnological studies from
the Antarctic and subantarctic islands. In: Pienitz, R., Douglas, S. V., Smol, J. P. [eds.]:
Long-term Environmental Change in Arctic and Antarctic Lakes. Springer, Dordrecht, s.
419–474.
Hodgson, D. A., Smol, J. P. (2008): High latitude paleolimnology. In: Vincent, W. F.,
Laybourn-Parry, J. [eds.]: Polar lakes and rivers – limnology of Arctic and Antarctic
aquatic ecosystems. Oxford University Press, Oxford, s. 43–64.
Hogg, A. G., Hua, Q., Blackwell, P. G., Niu, M., Buck, C. E., Guilderson, T. P., Heaton, T.
J., Palmer, J. G., Reimer, P. J., Reimer, R. W., Turney, C. S. M., Zimmerman, S. R. J.
(2013): SHCal13 southern hemisphere calibration, 0–50,000 cal BP. Radiocarbon, 55, č.
4, s. 1889–1903.
Hošek, J., Hambach, U., Lisá, L., Grygar, T. M., Horáček, I., Meszner, S., Knésl, I. (2015):
An integrated rock-magnetic and geochemical approach to loess/paleosol sequences from
78
Bohemia and Moravia (Czech Republic): Implications for the Upper Pleistocene
paleoenvironment in central Europe. Palaeogeography, Palaeoclimatology,
Palaeoecology, 418, č. 2, s. 344–358.
Howe, J. A., Pudsey, C. J. (1999): Antarctic circumpolar deep water: a Quaternary paleoflow
record from the northern Scotia Sea, South Atlantic ocean. Journal of Sedimentary
Research, 69, č. 4, s. 847–861.
Hrbáček, F., Nývlt, D., Láska, K. (2017): Active layer thermal dynamics at two lithologically
different sites on James Ross Island, Eastern Antarctic Peninsula. Catena, 149, č. 4, s.
592–602.
Hua, Q., Barbetii, M., Rakowski, A. Z. (2013): Atmospheric radiocarbon for the period
1950–2010. Radiocarbon, 55, č. 4, s. 2059–2072.
Humlum, O., Elberling, B., Hormes, A., Fjordheim, K., Hansen, O. H., Heinemeier, J.
(2005): Late-Holocene glacier growth in Svalbard, documented by subglacial relict
vegetation and living soil microbes. The Holocene, 15, č. 3, s. 396–407.
Ingólfsson, O., Hjort, Ch., Björck, S., Smith, L. R. I. (1992): Late Pleistocene and Holocene
glacial history of James Ross Island, Antarctic Peninsula. Boreas, 21, č. 3, s. 209–222.
Isaksson, E., Kohler, J., Pohjola, V., Moore, J., Igarashi, M., Karlöf, L., Martma, T., Meijer,
H., Motoyama, H., Vaikmäe, R., van de Wal, R. S. W. (2005): Two ice-core δ18
O records
from Svalbard illustrating climate and sea-ice variability over the last 400 years. The
Holocene, 15, č. 4, s. 501–509.
Isaksson, E., Pohjola, V., Jauhiainen, T., Moore, J., Pinglot, J. F., Vaikmäe, R., van de Wal,
R. S. W., Hagen, J. O., Ivask, J., Karlöf, L., Martma, T., Meijer, H. A. J., Mulvaney, R.,
Thomassen, M., van den Broeke, M. (2001): A new ice-core record from
Lomonosovfonna, Svalbard: viewing the 1920 – 97 data in relation to present climate and
environmental conditions. Journal of Glaciology, 47, č. 157, s. 335– 345.
Jones, V. J., Hodgson, D. A., Chepstow-Lusty, A. (2000): Palaeolimnological evidence for
marked Holocene environmental changes on Signy Island, Antarctica. The Holocene 10,
č. 1, s. 43–60.
Juggins, S. (2017): rioja: Analysis of Quaternary Science Data, R package version (0.9-15.1)
Kadey, F. L. (1983): Diatomite. In: LeFond, S. J. [eds.]: Industrial Rocks and Minerals.
American Institute of Mining, Metallurgical, and Petroleum Engineers, New York, s.
677–708.
Kaufman, D. S., Schneider, D. P., McKay, N. P., Ammann, C. M., Bradley, R. S., Briffa, K.
R., Miller, G. H., Otto-Bliesner, B. L., Overpeck, J. T., Vinther, B. M., Arctic Lakes 2k
Project Members (2009): Recent Warming Reverses Long-Term Arctic Cooling. Science,
325, č. 5945, s. 1236–1239.
79
King, J. C. (1994): Recent climate variability in the vicinity of the Antarctic Peninsula.
International Journal of Climatology, 14, č. 4, s. 357–369.
Klinck, J. M., Hofmann, E. E., Beardsley, R. C., Salihoglu, B., Howard, S. (2004): Watermass
properties and circulation on the west Antarctic Peninsula Continental Shelf in
Austral Fall and Winter 2001. Deep-Sea Research Part II – Topical Studies in
Oceanography, 51, č. 17–19, s. 1925–1946.
Koinig, K. A., Shotyk, W., Lotter, A. F., Ohlendorf, Ch., Sturm, M. (2003): 9000 years of
geochemical evolution of lithogenic major and trace elements in the sediment of an alpine
lake – the role of climate, vegetation, and land-use history. Journal of Paleolimnology,
30, č. 3, s. 307–320.
Kylander, M. E., Klaminder, J., Wohlfarth, B., Löwemark, L. (2013): Geochemical
responses to paleoclimatic changes in southern Sweden since the late glacial: the
Hässeldata Port lake sediment record. Journal of Paleolimnology, 50, č. 1, s. 57–70.
Larsen, D. J., Miller, G. H., Geirsdóttir, Á. (2013): Asynchronous little ice age glacier
fluctuations in Iceland and European alps linked to shifts in subpolar North Atlantic
circulation. Earth and Planetary Science Letters, 380, č. 20, s. 52–59.
Láska, K., Witoszová, D., Prošek, P. (2012): Weather patterns of the coastal zone of
Petuniabukta, central Spitsbergen in the period 2008–2010. Polish Polar Research, 33, č.
4, s. 309–326.
Last, W. M., Smol, J. P. (2001): Tracking Environmental Change Using Lake Sediments.
Volume 1: Basin Analysis, Coring, and Chronological Techniques. Kluwer Academic
Publishers, Dordrecht, The Netherlands, 573 s.
Leloup, M., Nicolau, R., Pallier, V., Yéprémian, C., Feuillade-Cathalifaud, G. (2013):
Organic matter produced by algae and cyanobacteria: Quantitative and qualitative
characterization. Journal of Environmental Sciences, 25, č. 6, 1089–1097.
Leng, M. J., Wagner, B., Anderson, N. J., Bennike, O., Woodley, E., Kemp, S. J. (2012):
Deglaciation and catchment ontogeny in coastal south-west Greenland: implications for
terrestrial and aquatic carbon cycling. Journal of Quaternary Science, 27, č. 6, s. 575–
584.
Leventer, A., Domack, E. W., Ishman, S. E., Brachfeld, S., McClennen, C. E., Manley, P.
(1996): Productivity cycles of 200–300 years in the Antarctic Peninsula region:
understanding linkages among the sun, atmosphere, oceans, sea ice, and biota. Geological
Society of America Bulletin, 108, č. 12, s. 1626–1644.
Lopez, P., Navarro, E., Marce, R., Ordoñez, J., Caputo, L., Armengol, J. (2006): Elemental
ratios in sediments as indicators of ecological processes in Spanish reservoirs. Limnetica,
25, č. 1–2, s. 499–512.
80
Maher, B. A., Thompson, R. (1999): Quaternary climates, environments and magnetism.
Cambridge University Press, Cambridge, 403 s.
Małecki, J. (2016): Accelerating retreat and high-elevation thinning of glaciers in central
Spitsbergen. The Cryosphere, 10, č. 3, s. 1317–1329.
Mann, M., Zhang, Z., Hughes, M. K., Bradley, R. S., Miller, S. K., Rutherford, S., Ni, F.
(2008): Proxy-based reconstructions of hemispheric and global surface temperature
variations over the past two millennia. Proceedings of the National Academy of Sciences,
105, č. 36, s. 13252–13257.
Mann, S., Sparks, N. H. C., Frankel, R. B., Bazylinski, D. A., Jannasch, H. W. (1990):
Biomineralization of ferrimagnetic greigite (Fe3S4) and iron pyrite (FeS2) in a
magnetotactic bacterium. Nature, 343, č. 6255, s. 258–261.
Marinsek, S., Ermolin, E. (2015): 10-year mass balance by glaciological and geodetic
methods of Glaciar Bahía del Diablo, Vega Island, Antarctic Peninsula. Annals of
Glaciology, 56, č. 70, s. 141–146.
Masson-Delmotte, V., Buiron, D., Ekaykin, A., Frezzotti, M., Gallée, H., Jouzel, J., Krinner,
G., Landais, A., Motoyama, H., Oerter, H., Pol, K., Pollard, D., Ritz, C., Schlosser, E.,
Sime, L.C., Sodemann, H., Stenni, B., Uemura, R., Vimeux, F. (2011): A comparison of
the present and last interglacial periods in six Antarctic ice cores. Climate of the Past, 7,
č. 2, s. 397–423.
Matsuoka, K. H., Narita, K., Sugiyama, S., Matoba, S., Motoyama, Watanabe, O. (1997):
Characteristics of AC-ECM signals obtained by use of the Vestfonna ice core, Svalbard.
Proceedings NIPR Symp. Polar Meteorolog. Glaciol. 11, č. 1, s. 67–76.
Minzoni, R. T., Anderson, J. B., Fernandez, R., Wellner, J. S. (2015): Marine record of
Holocene climate, ocean, and cryosphere interactions: Herber Sound, James Ross Island,
Antarctica. Quaternary Science Reviews, 129, č. 23, s. 239–259.
Moholdt, G., Nuth, C., Hagen, J. O., Kohler, J. (2010): Recent elevation changes of Svalbard
glaciers derived from ICESat laser altimetry. Remote Sensing of Environment, 114, č.
11, s. 2756–2767.
Morris, E. M., Vaughan, D. G. (2003): Spatial and temporal variations of surface
temperature on the Antarctic Peninsula and the limit of variability of ice shelves. In:
Domack, E. W., Leventer, A., Burnett, A., Bindschadler, R., Convey, P., Kirby, M. [eds.]:
Antarctic Peninsula Climate Variability: Historical and Paleoenvironmental Perspectives.
American Geophysical Union, Washington, D.C., s. 61–68.
Müller, J., Werner, K., Stein, R., Fahl, K., Moros, M., Jansen, E. (2012): Holocene cooling
culminates in sea ice oscillations in Fram Strait. Quaternary Science Reviews, 47, č. 12,
s. 1–14.
81
Mulvaney, R., Abram, N.J., Hindmarsh, R.C.A., Arrowsmith, C., Fleet, L., Triest, J., Sime,
L. C., Alemany, O., Foord, S. (2012): Recent Antarctic Peninsula warming relative to
Holocene climate and ice-shelf history. Nature, 489, č. 7414, s. 141–144.
Navas, A., López-Martínez, J., Casas, J., Machín, J., Durán, J.J., Serrano, E., Cuchi, J.-A.,
Mink, S. (2008): Soil characteristics on varying lithological substrates in the South
Shetland Islands, maritime Antarctica. Geoderma, 144, č. 1–2, s. 123–139.
Naveira Garabota, A. C., Heywood, K. J., Stevens, D. P. (2002): Modification and pathways
of Southern Ocean Deep Waters in the Scotia Sea. Deep-Sea Research I, 49, č. 4, s. 681–
705.
Nedbalová, L., Nývlt, D., Lirio, J. M., Kavan, J., Elster, J. (2017): Current distribution of
Branchinecta gaini on James Ross Island and Vega Island. Antarctic Science 29, č. 4, s.
341–342.
Nydal, R., Löwseth, K. (1983): Tracing bomb 14
C in the atmosphere 1962–1980. Journal of
Geophysical Research: Ocean, 88, č. C6, s. 3621–3642.
O’Dowd, C. D., Smith, M. H., Consterdine, I. E., Lowr, J. A. (1997): Marine aerosol, seasalt,
and the marine sulphur cycle: A short review. Atmospheric Environment, 31, č. 1,
73–80.
O’Sullivan, P. et al. (2008): The lakes handbook: limnology and limnetic ecology (Vol. 1).
John Wiley & Sons, USA, 709 s.
Oliva, M., Antonieades, D., Giralt, S., Granados, I., Pla Rabes, S., Toro, M., Sanjurjo, J.
(2016): La deglaciación de las áreas libres de hielo de las islas Shetland del Sur
(Antártida). Ejemplos de Byers (Livingston) y Barton (King George). Cuaternario y
Geomorfología, 30, č. 1–2, s. 105–118.
Orsi, A. H., Whitworth, T., III, Nowlin, W. D. Jr (1995): On the meridional extent and fronts
of the Antarctic Circumpolar Current. Deep-Sea Research I, 42, č. 5, s. 641–673.
Peel, D. A., Mulvaney, R. (1992): Time trends in the pattern of Ocean–atmosphere exchange
in an ice core from the Weddell Sea sector of Antarctica. Tellus Series B – Chemical and
Physical Meteorology, 4, č. 4, s. 430–442.
Pfeffer, M., Langmann, B., Graf, H.-F. (2006): Atmospheric transport and deposition of
Indonesian volcanic emissions. Atmospheric Chemistry and Physics, 6, č. 9, s. 2525–
2537.
Phillips, W., Briner, J. P., Gislefoss, L., Linge, H., Koffman, T., Fabel, D., Xu, S., Hormes,
A. (2017): Late Holocene glacier activity at inner Hornsund and Scottbreen, southern
Svalbard. Journal of Quaternary Science, 32, č. 4, 501–515.
Pielou, E. (1966): The measurement of diversity in different types of biological collections.
Journal of Theoretical Biology, 13, č. 10, s. 31–144.
82
Pinsel, E. (2014): Environmental Changes in a High Arctic Ecosystem. Diplomová práce.
Antwerpen, 160 s.
Pudsey, C. J., Evans, J. (2001): First survey of Antarctic sub-ice shelf sediments reveals
Mid-Holocene ice shelf retreat. Geology, 29, č. 9, s. 787–790.
Rabassa, J. (1983): Stratigraphy of the glacigenic deposits in northern James Ross Island,
Antarctic Peninsula. In: Evenson, E., Schlüchter, C., Rabassa, J. [eds.]: Tills and related
deposits. A. A. Balkema, Rotterdam, s. 329–40.
Rachlewicz, G., Szczucínski, W., Ewertowski, M. (2007): Post-“Little Ice Age” retreat rates
of glaciers around Billefjorden in central Spitsbergen, Svalbard. Polish Polar Research,
28, č. 3, s. 159–186.
Rasmussen, T. L., Thomsen, E., Skirbekk, K., Ślubowska-Woldengen, M., Klitgaard
Kristensen, D., Koç, N. (2014): Spatial and temporal distribution of Holocene
temperature maxima in the northern Nordic seas: interplay of Atlantic-, Arctic- and polar
water masses. Quaternary Science Reviews, 92, č. 9, s. 280–291.
Reimer, P. J., Bard, E., Bayliss, A., Beck, J. W., Blackwell, P. G., Bronk Ramsey, C., Buck,
C. E., Edwards, R. L., Friedrich, M., Grootes, P. M., Guilderson, T. P., Haflidason, H.,
Hajdas, I., Hatté, C., Heaton, T. J., Hoffmann, D. L., Hogg, A. G., Hughen, K. A., Kaiser,
K. F., Kromer, B., Manning, S. W., Niu, M., Reimer, R. W., Richards, D. A., Scott, E.
M., Southon, J. R., Turney, C. S. M., Van Der Plicht, J. (2013): IntCal13 and Marine13
radiocarbon age calibration curves, 0–50,000 years cal BP. Radiocarbon, 55, č. 4, s. 1869–
1887.
Reimer, P. J., Brown, T. A., Reimer, R. W. (2004): Discussion: reporting and calibration of
post-bomb 14
C data. Radiocarbon, 46, č. 3, s. 1299–1304.
Reynolds, J. M. (1981): The distribution of mean annual temperatures in the Antarctic
Peninsula. British Antarctic Survey Bulletin, 54, č. 1, s. 123–133.
Roberts, S. J., Hodgson, D. A., Bentley, M. J., Smith, J. A., Millar, I., Olive, V., Sugden, D.
E. (2008): The Holocene history of George VI Ice Shelf, Antarctic Peninsula from clastprovenance
analysis of epishelf lake sediments. Palaeogeography, Palaeoclimatology,
Palaeoecology, 259, č. 2–3, s. 258–283.
Roberts, E. M., Lamanna, M. C., Clarke, J. A., Meng, J., Gorscak, E., Sertich, J. J. W.,
O’Connor, P. M., Claeson, K. M., MacPhee, R. D. E. (2014): Stratigraphy and vertebrate
paleoecology of Upper Cretaceous – lowest Paleogene strata on Vega Island, Antarctica.
Palaeogeography, Palaeoclimatology, Palaeoecology, 402, č. 10, s. 55–72.
Robinson, S. A., Wasley, J., Tobin, A. K. (2003): Living on the edge–plants and global
change in continental and maritime Antarctica. Global Change Biology, 9, č. 12, s. 1681–
1717.
Roman, M. (2014): Holocenní vývoj arktických jezer. Bakalářská práce, Praha, 66 s.
83
Roman, M. (2017): Correlation of abiotic proxies in Holocene lacustrine sediments of PeriAtlantic
Arctic. Diplomová práce. Praha, 119 s.
Ross, R., Cox, E. J., Karayeva, N. I., Mann, D. G., Paddock, T. B. B., Simonsen, R., Sims,
P. A. (1979): An amended terminology for the siliceous components of the diatom cell.
Nova Hedwigia, 30, č. 1–4, s. 513–533.
Rothwell, R. G., Rack, F. R. (2006): New techniques in sediment core analysis: an
introduction. In: Rothwell, R. G. [eds.]: New Techniques in Sediment Core Analysis.
Geological Society, London, s. 1–29.
Røthe, T. O., Bakke, J., Vasskog, K., Gjerde, M., D’Andrea, W. J., Bradley, R. S. (2015):
Arctic Holocene glacier fluctuations reconstructed from lake sediments at Mitrahalvøya,
Spitsbergen. Quaternary Science Reviews, 109, č. 3, s. 111–125.
Round, F. E., Crawford, R. M., Mann, D. G. (1990): The Diatoms: biology and morphology
of the genera. Cambridge University Press, Cambridge, 760 s.
Salvigsen, O., Lauritzen, Ø., Mangerud, J. (1983): Karst and karstification in gypsiferous
beds in Mathiesondalen, Central Spitsbergen, Svalbard. Polar Research, 1, č. 1, s. 83–88.
Sarnthein, M., van Kreveld, S., Erlenkeuser, H., Grootes, P. M., Kucera, M., Pflaumann, U.,
Schulz, M. (2003): Centennial-to-millennial scale periodicities of Holocene climate and
sediment injections off the western Barents shelf, 75°N. Boreas, 32, č. 3, s. 447–461.
Shannon, C. E., Weaver W. (1949): The mathematical theory of communication. Urbana
University Press, Illinois, 125 s.
Schumacher, B. A. (2002): Methods for the Determination of Total Organic Carbon (TOC)
in Soils and Sediments. Ecological Risk Assessment Support Center, Office of Research
and Development, US Environmental Protection Agency, Las Vegas, 1, č. 1, s. 1–23.
Schwerdtfeger, W. (1975): The effect of the Antarctic Peninsula on the temperature regime
of the Weddell Sea. Monthly Weather Review, 103, č. 1, s. 45–51.
Simmonds, I. (2003): Modes of atmospheric variability over the Southern Ocean. Journal of
Geophysical Research – Oceans, 108, č. C4, s. 78–80.
Sjunneskog, C., Taylor, F. (2002): Postglacial marine diatom record of the Palmer Deep,
Antarctic Peninsula (ODPL Leg 178, Site 1098) 1. Total diatom abundance.
Paleoceanography and Paleoclimatology, 17, č. 3, s. 1–8.
Smellie, J. L., Johnson, J. S., McIntosh, W. C., Esser, R., Gudmundsson, M. T., Hambrey,
M. J., De Vries, B. V. W. (2008): Six million years of glacial history recorded in volcanic
lithofacies of the James Ross Island Volcanic Group, Antarctic Peninsula.
Palaeogeography, Palaeoclimatology, Palaeoecology, 260, č. 1–2, s. 122–148.
84
Smith, D. A., Hofmann, E. E., Klinck, J. M., Lascara, C. M. (1999): Hydrography and
circulation of the west Antarctic Peninsula continental shelf. Deep-Sea Research Part I –
Oceanographic Research Papers, 46, č. 6., s. 925–949.
Smith, D. A., Klinck, J. M. (2002): Water properties on the west Antarctic Peninsula
continental shelf: a model study of surface fluxes and sea-ice. Deep-Sea Research Part II
– Topical Studies in Oceanography, 49, č. 21, s. 4863–4886.
Solomina, O. N., Bradley, R. S., Hodgson, D. A., Ivy-Ochs, S., Jomelli, V., Mackintosh, A.
N., Nesje, A., Owen, L. A., Wanner, H., Wiles, G. C., Young, N. E. (2015): Holocene
glacier fluctuations. Quaternary Science Reviews, 111, č. 5, s. 9–34.
Spielhagen, R. F., Werner, K., Sørensen, S. A., Zamelczyk, K., Kandiano, E., Budeus, G.,
Husum, K., Marchitto, T. M., Hald, M. (2011): Enhanced modern heat transfer to the
Arctic by warm Atlantic water. Science, 331, č. 6016, s. 450–453.
Stensrud, D. J. (1996): Importance of low-level jets to climate: a review. Journal of Climate,
9, č. 8, s. 1698–1711.
Sterken, M., Roberts, S. J., Hodgson, D. A., Vyverman, W., Balbo, A. L., Sabbe, K.,
Moreton, S. G., Verleyen, E. (2012): Holocene glacial and climate history of Prince
Gustav Channel, northeastern Antarctic Peninsula. Quaternary Science Reviews, 31, č. 1,
s. 93–111.
Suess, H. E. (1955): Radiocarbon concentration in modern wood. Science, 122, č. 3166, s.
415–417.
Sugden, D. E., Bentley, M. J., Ó Cofaigh, O. (2006): Geological and geomorphological
insights into Antarctic ice sheet evolution. Philosophical Transactions of the Royal
Society A, 364, č. 1605–1625.
Sund, M., Eiken, T., Hagen, J. O., Kääb, A. (2009): Svalbard surge dynamics derived from
geometric changes. Annals of Glaciology, 50, č. 52, s. 50–60.
Svendsen, J. I., Mangerud, J. (1997): Holocene glacial and climatic variations on
Spitsbergen, Svalbard. The Holocene, 7, č. 1, s. 45–57.
Taboada, T., Cortizas, A. M., García, C., García-Rodeja, E. (2006): Particle-size
fractionation of titanium and zirconium during weathering and pedogenesis of granitic
rocks in NW Spain. Geoderma, 131, č. 1–2, s. 218–236.
Tavernier, I., Verleyen, E., Hodgson, D. A., Heirman, K., Roberts, S. J., Imura, S., Kudoh,
S., Sabbe, K., De Batist, M., Vyverman, W. (2014): Absence of Medieval Climate
Anomaly, Little Ice Age and twentieth century warming in Skarvsnes, Lutzow Holm Bay,
East Antarctica. Antarctic Science, 26, č. 5, s. 585–598.
Thomas, E. R., Dennis, P. F., Bracegirdle, T. J., Franzke, C. (2009): Ice core evidence for
significant 100-year regional warming on the Antarctic Peninsula. Geophysical Research
Letters, 36, č. 20, s. 1–5.
85
Thompson, L. G., Peel, D. A., Mosley-Thompson, E., Mulvaney, R., Dai, J., Lin, P. N.,
Davis, M. E., Raymond, C. F. (1994): Climate since 1510 AD on Dyer Plateau, Antarctic
Peninsula: Evidence for recent climate change. Annals of Glaciology, 20, č. 2, s. 420–
426.
Thompson, R., Clark, R. M., Boulton, G. S. (2012): Core Correlation. In: Birks, H. J. B.
[eds.]: Tracking Environmental Change Using Lake Sediments. Vol 5: Data Handling and
Numerical Techniques. Springer, Dordrecht. s. 415–430.
Thompson, R., Oldfield, F. (1986): Environmental Magnetism. George Allen and Unwin,
London, 227 s.
Tjallingii, R., Röhl, U., Kölling, M., Bickert, T. (2007): Influence of the water content on Xray
fluorescence core-scanning measurements in soft marine sediments. Geochemistry,
Geophysics, Geosystems, 8, č. 2, s. 1–12.
Vaasma, T. (2008): Grain-size analysis of lacustrine sediments: a comparison of pretreatment
methods. Estonian Journal of Ecology, 57, č. 4, s. 231–243.
Váczi, P., Barták, M., Nedbalová, L., Elster, J. (2011): Comparative analysis of temperature
courses in Antarctic lakes of different morphology: Study from James Ross Island,
Antarctica. Czech Polar Reports, 1, č. 2, s. 78–87.
Van Daele, M., Moernaut, J., Silversmit, G., Schmidt, S., Fontijn, K., Heirman, K.,
Vandoorne, W., De Clercq, M., Van Acker, J., Wolff, C., Pino, M., Urrutia, R., Roberts,
S.J., Vincze, L., De Batist, M. (2014): The 600 yr eruptive history of Villarrica Volcano
(Chile) revealed by annually laminated lake sediments. Bulletin of the Geological Society
of America 126, č. 3–4, s. 481–498.
Van der Bilt, W. G. M., Bakke, J., D’Andrea, W. J., Vasskog, K., Bradley, R. S., Ólafsdóttir,
S. (2015): Reconstruction of glacier variability from lake sediments reveals dynamic
Holocene climate in Svalbard. Quaternary Science Reviews, 126, č. 9, s. 201–218.
Van der Bilt, W. G. M., D’Andrea, W. J., Bakke, J., Balascio, N. L., Werner, J. P., Gjerde,
M., Bradley, R. S. (2018): Alkelone-based reconstructions reveal four-phase Holocene
temperature evolution for High Arctic Svalbard. Quaternary Science Reviews 183, č. 5,
s. 204–213.
Van der Werff, A. (1955): A new method of concentrating and cleaning diatoms and other
organisms. Verhandlungen Internationalen Vereinigung für Theoretische und
Angewandte Limnologie 2, 13, č. 1, s. 276–277.
Vaughan, D. G., Marshall, G., Connolley, W. M., Parkinson, C., Mulvaney, R., Hodgson,
D. A., King, J. C., Pudsey, C. J., Turner, J., Wolff, E. (2003): Recent rapid regional
climate warming on the Antarctic Peninsula. Climatic Change, 60, č. 3, s. 243–274.
Verleyen, E., Hodgson, D. A., Sabbe, K., Cremer, H., Emslie, S. D., Gibson, J., Hall, B.,
Satoshi, I., Kudoh, S., Marshall, G. J., McMinn, A., Melles, M., Newman, L., Roberts,
86
D., Roberts, S., Singh, S. M., Sterken, M., Tavernier, I., Verkulich, S., Van de Vyver, E.,
Van Nieuwenhuyze, W., Wagner, B., Vyverman, W. (2011): Postglacial climate
variability along the east Antarctic coastal margin evidence from shallow marine and
coastal terrestrial records. Earth-Science Reviews, 104, č. 4, s. 199–212.
Vincke, S., Van de Vijver, B., Gremmen, N., Beyens, L. (2006): The moss dwelling
testacean fauna of the Stromness Bay (South Georgia). Acta Protozoologica, 45, č. 1–4,
65–75.
Walkley, A., Black, I. A. (1934): An examination of the Degtjareff method for determining
soil organic matter, and a proposed modification of the chromic acid titration method.
Soil Science, 37, č. 1, s. 29–38.
Wellner, J. S., Heroy, D. C., Anderson, J. B. (2006): The death mask of the Antarctic Ice
Sheet: comparison of glacial geomorphic features across the continental shelf.
Geomorphology, vol. 75, č. 1–2, s. 157–171.
Weltje, G. J., Tjallingii, R. (2008): Calibration of XRF core scanners for quantitative
geochemical logging of sediment cores: theory and application. Earth and Planetary
Science Letters, 274, č. 3–4, s. 423–438.
Werner, A. (1993): Holocene moraine chronology, Spitsbergen, Svalbard: Lichenometric
evidence for multiple Neoglacial advances in the Arctic. The Holocene, 3, č. 2, s. 128–
137.
Werner, K., Müller, J., Husum, K., Spielhagen, R. F., Kandiano, E. S., Polyak, L. (2016):
Holocene sea subsurface and surface water masses in the Fram Strait – Comparisons of
temperature and sea-ice reconstructions. Quaternary Science Reviews, 147, č. 17, s. 194–
209.
Wetzel, R. G. (2001): Limnology: Lake and River ecosystems. Academic Press, San Diego,
1006 s.
Wilkin, R. T., Barnes, H. L. (1997): Formation processes of framboidal pyrite. Geochimica
et Cosmochimica Acta, 61, č. 2, s. 323–339.
Winkelmann, D., Knies, J. (2005): Recent distribution and accumulation of organic carbon
on the continental margin west off Spitsbergen. Geochemistry, Geophysics, Geosystems,
6, č. 9, s. 1–22.
Wolfe, A. P., Miller, G. H., Olsen, C. A., Forman, S. L., Doran, P. T., Holmgren, S. U.
(2004): Geochronology of High Latitude lake sediments. In: Pienitz, R., Douglas, S. V.,
Smol, J. P. [eds.]: Long-term Environmental Change in Arctic and Antarctic Lakes.
Springer, Dordrecht, s. 419–474.
Worm, H. U. (1998): On the superparamagnetic – stable single domain transition for
magnetite, and frequency dependence of susceptibility. Geophysical Journal
International, 133, č. 1, s. 201–206.
87
Xu, H., Liu, B., Wu, F. (2010): Spatial and temporal variations of Rb/Sr ratios of the bulk
surface sediments in Lake Qinghai. Geochemical Transactions, 11, č. 1, s. 1–8.
Yoon, H. I., Park, P. K., Kim, Y., Kang, Yun, Cheon, Kang, S. H. (2003): Origins and
paleoceanographic significance of the layered diatom ooze interval from the Bransfield
Straight in the northern Antarctic Peninsula around 2500 years BP. In: Domack, E. W.,
Leventer, A., Burnett, A., Bindschadler, R., Convey, P., Kirby, M. [eds.]: Antarctic
Peninsula Climate Variability: Historical and Paleoenvironmental Perspectives.
American Geophysical Union, Washington, D.C., s. 225–238.
Zale, R., Karlén, W. (1989): Lake sediment cores from the Antarctic Peninsula and
surrounding island. Geografiska Annaler: Series A, Physical Geography, 71, č. 3–4, s.
211–220.
Zhang, Y., Renssen, H., Seppä, H. (2016): Effects of melting ice sheets and orbital forcing
on the early Holocene warming in the extratropical Northern Hemisphere. Climate of the
Past, 12, č. 5, s. 1119–1135.
Zidarova, R., Kopalová, K., Van de Vijver, B. (2016): Diatoms from the Antarctic Region.
I: Maritime Antarctica. Iconographia Diatomologica, vol. 24. Hardcover, Königstein, 509
s.
Maps and atlases
Dallmann, W. K., Piepjohn, K., Blomeier, D. (2004): Geological Map of Billfjorden, Central
Spitsbergen, Svalbard: With Geological Excursion Guide. 1:50 000. Norsk polarinstitutt,
Oslo.
Smellie, J. L., Johnson, J. S., Nelson, A. E. (2013): Geological map of James Ross Island. 1.
James Ross Island Volcanic Group. BAS GEOMAP 2 Series, Sheet 5. 1:125 000. British
Antarctic Survey, Cambridge, UK
88
APPENDIX
APPENDIX
Table 1 Clam’s (R) modelled b2k ages in the Lake Anonima. (Appendix 1)
depth (cm) min 95 % best-age max 95 %
0.0 –15 –13 –11
0.5 –10 –8 –6
1.0 –4 –3 –1
1.5 1 2 4
2.0 6 8 9
2.5 12 13 14
3.0 17 18 19
3.5 22 23 24
4.0 27 28 29
4.5 32 33 34
5.0 38 38 39
5.5 43 44 44
6.0 67 68 69
6.5 90 92 95
7.0 113 117 120
7.5 137 141 146
8.0 160 166 171
8.5 183 190 197
9.0 207 214 222
9.5 230 239 248
10.0 253 263 273
10.5 277 288 299
11.0 300 312 324
11.5 323 337 350
12.0 346 361 375
12.5 370 385 401
13.0 393 410 427
13.5 416 434 452
14.0 440 459 478
14.5 463 483 503
15.0 486 507 529
15.5 510 532 554
16.0 533 556 580
16.5 556 581 605
17.0 597 621 645
17.5 637 661 685
18.0 677 701 726
18.5 716 741 767
19.0 755 781 809
19.5 793 821 851
20.0 831 861 893
20.5 870 901 935
21.0 908 941 976
depth (cm) min 95 % best-age max 95 %
21.5 945 981 1018
22.0 983 1021 1060
22.5 1021 1061 1102
23.0 1058 1101 1144
23.5 1095 1142 1187
24.0 1134 1182 1229
24.5 1171 1222 1272
25.0 1209 1262 1314
25.5 1246 1302 1357
26.0 1284 1342 1400
26.5 1322 1382 1443
27.0 1359 1422 1486
27.5 1396 1462 1529
28.0 1433 1502 1572
28.5 1470 1542 1614
29.0 1482 1553 1625
29.5 1494 1564 1634
30.0 1506 1575 1644
30.5 1519 1586 1654
31.0 1531 1598 1664
31.5 1543 1609 1674
32.0 1555 1620 1685
32.5 1567 1631 1695
33.0 1579 1642 1706
33.5 1590 1653 1716
34.0 1602 1664 1726
34.5 1614 1675 1737
35.0 1625 1686 1747
35.5 1637 1697 1758
36.0 1648 1708 1769
36.5 1659 1719 1780
37.0 1671 1731 1792
37.5 1682 1742 1803
38.0 1694 1753 1814
38.5 1705 1764 1826
39.0 1717 1775 1837
39.5 1728 1786 1850
40.0 1739 1797 1861
40.5 1751 1808 1874
41.0 1762 1819 1886
41.5 1773 1830 1898
42.0 1784 1841 1910
42.5 1795 1852 1921
43.0 1806 1864 1932
43.5 1817 1875 1943
44.0 1828 1886 1955
depth (cm) min 95 % best-age max 95 %
44.5 1839 1897 1966
45.0 1850 1908 1977
45.5 1861 1919 1988
46.0 1871 1930 2000
46.5 1882 1941 2011
47.0 1893 1952 2022
47.5 1904 1963 2033
48.0 1915 1974 2045
48.5 1926 1985 2056
49.0 1936 1996 2068
49.5 1947 2008 2080
50.0 1958 2019 2092
50.5 1968 2030 2104
51.0 1979 2041 2115
51.5 1990 2052 2127
52.0 2000 2063 2139
52.5 2011 2074 2151
53.0 2021 2085 2164
53.5 2031 2096 2176
54.0 2041 2107 2188
54.5 2051 2118 2200
55.0 2061 2129 2213
55.5 2070 2141 2225
56.0 2080 2152 2238
56.5 2089 2163 2251
57.0 2099 2174 2265
57.5 2109 2185 2278
58.0 2118 2196 2292
58.5 2127 2207 2305
59.0 2136 2218 2319
59.5 2145 2229 2333
60.0 2154 2240 2346
60.5 2164 2251 2360
61.0 2173 2262 2373
61.5 2183 2274 2387
62.0 2191 2285 2401
62.5 2200 2296 2415
63.0 2209 2307 2429
63.5 2218 2318 2443
64.0 2227 2329 2456
64.5 2237 2340 2469
65.0 2246 2351 2483
65.5 2255 2362 2497
66.0 2264 2373 2511
66.5 2272 2384 2525
67.0 2281 2395 2538
depth (cm) min 95 % best-age max 95 %
67.5 2290 2407 2552
68.0 2299 2418 2566
68.5 2308 2429 2580
69.0 2317 2440 2593
Table 2 Clam’s (R) modelled b2k ages in the Lake Mathiesondalen 3. (Appendix 1)
depth (cm) min 99 % best-age max 99 %
0.00 –21 –18 –14
0.25 –19 –16 –13
0.50 –16 –13 –11
0.75 –14 –11 –8
1.00 –11 –8 –6
1.25 –9 –6 –3
1.50 –5 –2 1
1.75 –1 2 5
2.00 3 6 10
2.25 8 12 15
2.50 14 18 22
2.75 20 25 29
3.00 28 33 37
3.25 35 41 47
3.50 43 50 57
3.75 51 59 67
4.00 59 69 78
4.25 67 79 90
4.50 75 89 103
4.75 83 99 115
5.00 92 110 129
5.25 101 122 143
5.50 109 133 157
5.75 118 145 172
6.00 126 157 187
6.25 136 169 202
6.50 145 181 218
6.75 154 194 234
7.00 163 207 250
7.25 172 219 267
7.50 182 233 283
7.75 192 246 300
8.00 202 259 317
8.25 212 273 334
8.50 221 286 351
8.75 231 300 368
9.00 240 313 386
9.25 249 326 403
depth (cm) min 99 % best-age max 99 %
9.50 259 340 420
9.75 269 353 437
10.00 278 366 454
10.25 288 379 471
10.50 297 392 488
10.75 306 405 504
11.00 316 418 520
11.25 325 431 536
11.50 334 443 552
11.75 343 455 567
12.00 352 467 582
12.25 361 479 597
12.50 370 490 611
12.75 379 502 624
13.00 387 512 638
13.25 395 523 650
13.50 403 533 663
13.75 412 543 674
14.00 420 553 685
14.25 428 562 696
14.50 436 571 705
14.75 443 579 714
15.00 451 587 723
15.25 457 594 730
15.50 464 601 737
15.75 471 607 743
16.00 477 613 748
16.25 483 618 753
16.5 489 623 756
16.75 495 627 760
17.00 500 632 763
17.25 506 636 767
17.50 511 640 769
17.75 516 645 773
18.00 521 649 777
18.25 526 653 781
18.50 530 658 786
18.75 535 663 791
19.00 539 667 796
19.25 543 672 801
19.50 546 676 805
19.75 550 679 809
20.00 552 682 812
20.25 554 684 815
20.50 556 687 819
20.75 558 689 820
depth (cm) min 99 % best-age max 99 %
21.00 560 692 823
21.25 563 694 825
21.50 565 696 827
21.75 567 698 829
22.00 568 699 831
22.25 569 701 833
22.50 569 702 835
22.75 569 703 837
23.00 570 704 839
23.25 566 703 841
23.50 557 700 842
Table 3 Correlation (r) matrix of geochemical data in the Lake Anonima. (Appendix 2)
TOC LE Si Fe Ca Al Ti Mn Sr Rb Zr Ni
Ni -0.79 -0.84 0.39 0.88 0.83 0.68 0.80 0.86 0.87 -0.62 0.65
Zr -0.60 -0.65 0.40 0.67 0.59 0.53 0.54 0.69 0.69 -0.20
Rb 0.70 0.76 -0.29 -0.77 -0.82 -0.61 -0.71 -0.71 -0.77
Sr -0.93 -0.97 0.46 0.99 0.97 0.78 0.86 0.95
Mn -0.89 -0.93 0.47 0.95 0.93 0.75 0.82
Ti -0.84 -0.79 0.10 0.87 0.87 0.50
Al -0.70 -0.87 0.87 0.78 0.78
Ca -0.92 -0.97 0.46 0.97
Fe -0.92 -0.97 0.46
Si -0.34 -0.61
LE 0.90
TOC
Table 4 Correlation (r) matrix of geochemical data (dried samples) in the Lake Mathiesondalen 3. (Appendix
2)
TOC LE Si Fe Ca Al Ti Mn Sr Rb Zr Ni
Ni -0.56 0.09 0.70 0.50 -0.60 0.77 0.31 -0.50 -0.62 0.75 0.59
Zr -0.32 0.32 0.34 0.77 -0.55 0.45 0.30 -0.12 -0.49 0.78
Rb -0.46 0.36 0.51 0.86 -0.72 0.60 0.29 -0.17 -0.64
Sr 0.09 -0.22 -0.41 -0.36 0.47 -0.54 -0.13 0.60
Mn 0.13 0.11 -0.38 0.24 0.16 -0.42 -0.17
Ti -0.51 -0.16 0.52 0.20 -0.31 0.53
Al -0.70 -0.08 0.94 0.38 -0.63
Ca 0.51 -0.69 -0.69 -0.63
Fe -0.37 0.44 0.28
Si -0.69 -0.03
LE 0.03
TOC