Petrological and Geochronological investigation of the Lundy granite and its role in the
North Atlantic Igneous Province (NAIP)
High precision U-Pb dating and comparative study of Lundy granite and Fur Island tephra
Karlo Lisica
Thesis submitted for the degree of
Master of Science in Petrology and Geochemistry 60 credits
Centre for Earth Evolution and Dynamics Department of Geosciences
The Faculty of Mathematics and Natural Sciences UNIVERSITY OF OSLO
September 2021
© Karlo Lisica 2021
Petrological and Geochronological investigation of the Lundy granite and its role in the North Atlantic Igneous Province (NAIP)
Karlo Lisica
http://www.duo.uio.no/
Print: Reprosentralen, University of Oslo
I. ABSTRACT
To understand the consequences of current climate change, it is important to understand the changes in the carbon cycle and their timing to construct accurate models of climate change. For that reason, one of the best analogues, and most recent global warming event is the Palaeocene-Eocene Thermal Maximum, or PETM that started around 56 million years ago. One of the proposed theories of what caused PETM is the extensive volcanism linked to opening of the North Atlantic Igneous Province (NAIP). The timing and duration of the PETM is most notably recorded in the Danish island of Fur, where over 180 ash layers are preserved in a sedimentary sequence where the onset, main body and recovery of the PETM CIE can be observed. The recovery phase of PETM CIE is marked by thick horizon named ‘Ash -33’, a peraluminous rhyolite layer found in the sequence of mostly basaltic layers. To better constrain the resolution of timing and duration of the PETM, the magnitude of the activity in the NAIP during Paleogene, it is important to know the provenance of the ashes within the Fur Formation in Denmark. Lundy has been proposed to be volcanic source to the rhyolitic ash layers:
-33, +13, +19. The aims of this study are numerous: improving the temporal resolution of the Lundy magmatic complex, finding a source to the studied ash layers and to fill the gap on drivers of PETM.
This study utilizes state-of-the-art CA-ID-TIMS geochronology analysis and ICP-MS with goal to obtain precise and accurate data on the age and geochemistry coupled with the methods standard in petrology. The CA-ID-TIMS data yielded the age of 57.287 ± 0.076 Ma for the granite emplacement and the age of 55.938 ± 0.041 Ma trachyte emplacement, updating the previous data that reported early Paleocene ages. The Lundy rocks and selected ashes exhibit similarities in their major and trace- element geochemistry. Lundy granite and ash -33 are similar, with exception in the concentrations of HFSE. Lundy trachyte has noticeable similarities in both LILE and HFSE trends to ashes +13 and +19.
The current stratigraphy and the age presented in this study, although similar in geochemical trends, does not confirm for the main phase of the Lundy complex as the source of the ash -33. As the activity of the Lundy complex is not synchronous with PETM age model at Fur, it is unlikely that it caused long-term climate disturbances that are recorded in the Danish Fur Formation. Trachyte volcanic event is synchronous to the onset of PETM, marked by regional cooling effect that is observed bellow beginning of Fur Formation. Given the poor age constraint of the PETM model at Fur, it is unclear what is the exact relationship between activity of the Lundy complex and Fur Formation. Therefore, the detailed geochronological and geochemical study of Fur Formation can fill this gap.
II. PREFACE
My first introduction to igneous processes and volcanism was during the time I was writing my bachelor thesis: rift volcanism of East African Rift at University of Zagreb. The introduction to research and the topics I was particularly interested in, made me want to learn more. In December 2018, I contacted Morgan and Lars, now my supervisors, for my potential thesis topic. During our correspondence they introduced me to the how Fur ash formation and Lundy Island could be connected. Upon further discussion, I have found the interdisciplinary study of petrology, geochemistry, and (at the time) unknown to me, geochronology- challenge I was initially looking for.
Unfortunately, in the end of March of 2020, a new, unpredicted, situation appeared in a form of a global pandemic. This pandemic has not only caused a shift on global level but on individual one as well. For that reason, travelling became limited, and I have not been able to go to the scheduled field trip to Lundy. Luckily, this inability to travel has not affected the results in any way, other than samples arriving later than what was planned. Additionally, as the time was ticking, the research plan also changed. Although this slowed down the process of research, ultimately the thesis was written on time.
This thesis was written with an effort so everyone interested in the topics of petrology, geochemistry and geochronology can understand the goal of it, even if one has no geological background. For that reason, I have included as much figures as I could, all in hopes it would be easier to understand the interplay of different disciplines. Having said that, it is an advantage for a reader to be confident with geological concepts, especially as the data presented in this thesis is new and exciting, and understanding them will explain why this topic is important to research.
III. ACKNOWLEDGMENT
Anyone who has ever written a master thesis, will agree, is a difficult thing. The stress, lack of research experience and knowledge and general ‘pains’ of being an international student in a foreign country does not make this process any easier. Luckily, this thesis was not only a result of individual effort, but of group of people that were there to teach me and support me in my goal of becoming a scientist.
First, I would like to direct my utmost gratitude to my supervisors Morgan Jones and Lars Eivind Augland, for giving a chance and opportunity to an unknown from Croatia. Thank you, Morgan, for explaining over and over, what my thesis was about and overall kindness and support, especially with the difficulties during the pandemic. I am grateful that you were always encouraging and available for discussion and questions I had. Thank you, Lars, for explaining to me the workings of TIMS-lab and many possibilities of U-Pb dating methods in geoscience. I am grateful that you were always ready to help me, and for sparking the interest in this field!
Off-course, I must mention the army of people from labs at UiO and CEED (and outside of Norway!), that without them, this thesis would never be finished.
Gunborg Bye Fjeld for guidance on sample preparation. Our interest in unique outdoor activities and stories behind them made the time in lab fun. Salahalldin Akhavan for finalizing the thin sections and turning the blind eye to cutting my personal collection of rocks. Mufak Said Naoroz for guiding me through particle size analysis. Siri Simonsen for showing me the SEM. John Stevenson for helping me develop the ash transport model and understand the workings of the code. To employees of CEED for answering questions I had, that I felt were too embarrassing to ask my supervisors. Special gratitude to Dougal Jerram and Adam Beresford-Browne for collecting the samples in my name and subsequent zoom meetings. Your inputs with the thesis and the information on BPIP are acknowledged.
I am obliged to mention my friends from Oslo, who made this experience less stressful. Your patience with me and willingness to laugh at my stupid stories made me genuinely happy. I know our friendship will endure geographical borders!
This journey would never be possible with the unconditional love and support from my friends and family from Croatia. To prof. Petrinec, thank you for believing in me from the start. To the Danijel Popović fan club, I cherish the memories we have built together. Thank you, Sis, for being my biggest supporter and for explaining to the family why it is a good idea for me to move to Norway. Lastly, to my Balkan friends I met in Oslo, hvala vam na podršci i svim lijepim trenutcima u Oslu :)
TABLE OF CONTENTS
1. INTRODUCTION AND OBJECTIVES OF THE STUDY ... 1
2. SCIENTIFIC FRAMEWORK ... 5
The Paleocene- Eocene period: paleogeography and climate ... 5
Paleocene- Eocene Thermal Maximum (PETM): Timing, duration, causes & effects ... 7
The North Atlantic Igneous Province and British Paleogene Igneous Province: History, composition, and ages ... 10
Impact of NAIP: can volcanism cause climate change? ... 13
3. STUDY LOCATIONS ... 15
Geology of Lundy Island ... 15
Geology of Fur Island ... 18
4. THEORETICAL BACKGROUND AND INSTRUMENTATION ... 20
Geochemistry: trace elements as ‘fingerprints’ in geology ... 20
Isotopes, radioactivity, and decay... 21
Principles of U-Pb geochronology ... 22
Electron microbeams and mass spectrometry ... 25
5. METHODS AND MATERIALS ... 27
Field work ... 27
5.1.1 Sampling and rock description ... 27
Lab work ... 35
5.2.1 Thin sections for petrography ... 35
5.2.2 Zircon extraction ... 35
5.2.3 Scanning Electron Microscopy (SEM) ... 36
5.2.4 Geochronology: CA-ID-TIMS analysis ... 36
5.2.5 Preparation for major and trace geochemistry ... 37
5.2.6 Ash transport model: particle size analysis ... 37
6. RESULTS ... 38
Thin section analysis ... 39
6.1.1 Lundy Granite: ... 39
6.1.2 Lundy Dykes: ... 41
Zircon description and SEM analysis ... 43
6.2.1 OS-LD1 ... 44
6.2.2 OS-LD10 ... 45
Geochronology ... 46
6.3.1 Age of OS-LD 1 ... 48
6.3.2 Age of OS-LD 10 ... 49
Geochemistry ... 50
How windy/explosive does it need to be to transport ash grain to great distances? ... 53
7. DISCUSSION ... 55
Rock vs. zircon: the link between rock crystallization history and zircon texture ... 56
The Lundy Igneous Complex: the age of granite emplacement and the age of volcanic event ... 58
Lundy vs. Fur: comparison of geochemical trends ... 59
The origin of rhyolitic ash in Fur Formation: is Lundy the source? ... 60
Implications of ash model: could Lundy activity play a role in the climate change of Paleocene- Eocene? ... 62
8. CONCLUSIONS ... 64
Recommendations on further research ... 65
9. REFRENCE LIST ... 66
10. APPENDIX ... 66
Table with complete geochemistry data ... 73
Ash transport model: Full workbook ... 74
TABLE OF FIGURES AND TABLES
Figure 1.2: Compilation of data from the late Paleocene and early Eocene... 3
Figure 2.1: Chronostratigraphic chart of Paleogene period and isotope curve graph ... 6
Figure 2.2: Paleogeographic reconstruction of continents at different stages. ... 7
Figure 2.3: Emplacement methods of the NAIP and possible causes of climate change in the Paleogene. ... 8
Figure 2.4: Paleogeographic reconstruction of the North Atlantic Igneous Province at ~56 Ma ... 11
Figure 2.5: The compiled age map of British Paleogene Igneous Province and its main igneous intrusions. . 13
Figure 3.1: A geological map of Lundy ... 15
Figure 3.2: Distribution of dyke swarms and major faults in the greater area of BPIP ... 17
Figure 3.3: The elevation map and photo of the locality of Stolleklint ... 18
Figure 3.4: The detail of stratigraphy at Stolleklint ... 19
Figure 4.1: A diagram of compatible vs. incompatible elements ... 21
Figure 4.2: The distribution of natural stable isotopes in the neutron-proton diagram ... 22
Figure 4.3: The chart of nuclides showing the decay series of 232Th, 235U and 238U ... 23
Figure 4.4: Wetherill & Tera-Wasserburg concordia diagram ... 25
Figure 4.5: Schematic diagram of a sector-field noble gas or TIMS mass spectrometer ... 26
Figure 5.1: The field trip itinerary ... 28
Figure 5.2: The preparation process for extracting zircons from solid materials ... 35
Figure 6.1. Microscopic image of the CGMBG ... 39
Figure 6.2: Microscopic image of Biotite megacryst ... 40
Figure 6.3: Microscopic image of the M/FGBG ... 40
Figure 6.4: Microscopic image of the Basalt-diorite dyke ... 41
Figure 6.5: Microscopic image of the trachyte dyke ... 42
Figure 6.6: OS-LD 1 and OS-LD 10 zircon population ... 43
Figure 6.7: Cathodoluminescence (CL) images of zircons chosen for dating from the sample OS-LD1 ... 44
Figure 6.8: Cathodoluminescence (CL) images of zircons chosen for dating from the sample OS-LD10 ... 45
Figure 6.9: Concordia diagram for OS-LD1 ... 48
Figure 6.10: Concordia diagram for OS-LD10 ... 49
Figure 6.11: The plots of classification diagrams ... 51
Figure 6.12: Granite tectonic discrimination ... 52
Figure 6.13: Spider diagrams ... 53
Figure 6.14: The output of an ash transport modelling ... 54
Figure 7.1: Spider diagrams for selected samples ... 59
Figure 7.2: Correlation of Lundy activity and Fur Formation ... 61
Table 1: ALS Geochemistry codes, description, and instruments. ... 37
Table 2: U-Pb data for selected zircons from the samples OS-LD1 and OS-LD10 ... 47
Table 3: Major element (in wt %) and minor element (in ppm) data ... 50
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1. INTRODUCTION AND OBJECTIVES OF THE STUDY
One of the most important topics of the modern age is climate change. There are already many observable effects of global climate change on the environment: shrinking of glaciers, accelerated sea level rise, longer and more intense heat waves, storms, ocean warming, and acidification of oceans to name just a few. According to NASA the Earth’s average temperature has increased about 1°C since the start of the industrial era. Flora and fauna distributions will change, potentially causing extinction events. Although we as a species are adaptable, it is inevitable that climate change will affect us as well. Research on the current climate change involves creating climate models in hopes to predict the consequences of the anthropogenic carbon emissions.
What will happen to the environment with increasing temperature? How will climate change affect the human race? These are some of the questions posed by climate change and to scientists who are trying to find answers to this problem. To fully understand the consequences of current climate change, it is logical to look back into the past. In words of Charles Lyell, the past unravels processes that, by studying them, reveals what could happen in the future. Understanding the changes in the carbon cycle and their timing is important to construct accurate models of climate change. For that reason, one of the best analogues, and most recent global warming event is the Paleocene-Eocene Thermal Maximum, or PETM that started around 56 million years ago (Storey et al. 2007a; Svensen et al. 2019). This event was marked by the significant release of carbon into the atmosphere and oceans that caused many negative effects; some of which can be directly compared due to recent global warming (Figure 1.1).
Figure 1.1: Observable effects of global warming that is only the beginning of processes caused by rapid climate change.
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One of the main theories what have caused the PETM, is that it was a consequence of volcanic activity associated with the emplacement of the North Atlantic Igneous Province (NAIP). The theory that extensive volcanism can cause global climate changes and possibly trigger extinction events is widely recognized (review in Bond & Wignall, 2014). Koch et al. (1992) were the first to identify the PETM by measuring the negative carbon isotopic excursion (CIE) in pedogenic carbonate and mammalian tooth enamel derived from continental rocks. They recognized that the CIE coincides with the extinction of large mammals. Since then, recent studies (e.g. Jones et al. 2019; Stokke et al. 2020a) are working to increase the record resolution of PETM across the NAIP using different elements. The change in global surface temperatures is well preserved and recorded in the sediments all around NAIP.
The timing and duration of the PETM is most notably recorded in the Danish island of Fur,where over 180 ash layers are preserved in a sedimentary sequence where the onset, main body and recovery of the PETM CIE can be observed. Most of these ashes are tholeiitic basalts in composition. However, there are a couple of peraluminous silicic tephras, including the thick marker horizon named ‘Ash-33’
(see review in Larsen et al. 2003). This layer is found at the end CIE of the Palaeocene-Eocene Thermal Maximum (PETM) (Jones et al. 2019). It is proposed that the layer -33 may play an important role in the recovery phase, as the volcano-derived CO2 has sequestrated as cement in the tephras found across NAIP (Longman et al. 2021). The PETM is of interest because it also coincides with the peak activity of the NAIP (Wilkison et al. 2017). This temporal correlation (Figure 1.2) between the PETM and the NAIP suggests that the NAIP may be the cause of, or at least a catalyst to, rapid global warming, which is a hypothesis that that needs to be further explored.
The focus area of this thesis is Lundy, a small island within the Bristol Channel in the United Kingdom.
Lundy is the southernmost part of British Paleogene Igneous Province (BPIP), an off-axis part of the NAIP. In the BPIP, most of the magmatism is expressed as basaltic flood magmatism, with associated intrusive centres. Lundy has been a research interest for a long time. The first investigations aimed to constrain the petrology of the island (Dollar et al. 1941; Edmonds et al. 1979). Later, the research evolved into understanding the geochemistry and how mantle and tectonic processes shaped Lundy’s granite (Stone, 1990; Thorpe et al. 1990; Thorpe & Tindle, 1992). Recently, efforts have been made to understand the Lundy’s role in the BPIP, that is, its role in distal NAIP processes. Dating different dyke sets provides temporal constraints for pulses of extension and rifting, magma initiation, and the end of magmatism for volcanic products found across NAIP (e.g. BPIP: Hamilton et al. 1998;
Chambers et al. 2005; Ganerød et al. 2011; Charles et al. 2017).
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The granite body is anomalous, not only geographically (most southern expression of the BPIP and NAIP magmatism) but petrologically and geochemically as well (its S-type characteristics and geochemical affinities: peraluminous, high Al, high Rb, Rb/Sr and high trace alkali, Nb, U) (Thorpe et al. 1990; Thorpe & Tindle, 1992; Larsen et al. 2003; Charles et al. 2017). For this reason, Lundy has been proposed to be a source volcano for the ash layer -33, as well as ash layers +13 and +19 found across North Sea and Denmark.
Figure 1.2: Compilation of data from the late Paleocene and early Eocene. This work will try to correlate the Fur Formation (specifically ash layers -33, +13 and +19) with the NAIP activity in the area of BPIP (specifically Lundy Island). Since Lundy shares very similar geochemical signatures with the rhyolite ash layers, in order to correlate it with Fur and PETM event it should have had volcanic activity at the time of PETM (~56 Ma). The time scale is sourced from International Commission on Stratigraphy, 2021. Carbon and Oxygen isotope data from Litter et al. 2014, indicates hyperthermal events (PETM and ETM2). The overview of NAIP activity is a compilation of age data from Wilkinson et al. 2017.
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To better constrain the resolution of timing and duration of the PETM, the magnitude of the activity in the NAIP during the Paleogene, it is important to know the provenance that is the source(s) of the ashes within the Fur Formation in Denmark. To do so this thesis will investigate several hypotheses:
1) Could Lundy granite be a volcanic source for rhyolitic ash layers preserved in the Fur Formation in Denmark, particularly ash layers -33, +13 and +19? If so, Lundy could be a key in tracing the processes and the timing of the NAIP due to its specific chemistry and geographic position to the rest of BPIP;
2) What is the magnitude of the volcanic activity that formed the rhyolite ash layers? Given the thickness of the layers (on average 15cm at Fur), grain size and general morphology of grains, what are the necessary plume heights to transport ash at large distances (>1000km)?
3) Could activity of Lundy complex be connected to climate perturbations of Paleocene-Eocene?
Apart from main hypotheses, this study will explore the relation between zircon texture and magma evolution of each of the phases, using age data and the known isotopic data (Thorpe et al. 1990; Thrope
& Tindle, 1992).
To answer these questions the project plan is as follows:
1) Collecting the granite and different dyke samples from Lundy;
2) Do petrography to determine the mineralogy of the samples;
3) CL-SEM analysis to acquire internal textures of zircons from different rock phases.
4) Acquiring precise chronology of Lundy emplacement using state-of-the-art CA-ID-TIMS:
New and high-resolution dates using U-Pb method on selected zircons. The aim is to significantly improve the temporal resolution of the magmatic complex as compared to that presented in Charles et al. 2017;
5) Comparison of the major and minor element geochemistry of the Lundy rocks and geochemically similar ash layers at Fur, specifically -33, +13 and +19;
6) Designing a robust ash transport model using grain size data based on the work of Stevenson et al. 2015.
If done successfully this thesis will open new questions worth investigating but, more importantly, it will play a role in understanding the PETM-CIE age model, creating at the same time, an improved temporal framework for pulses of rifting, magma initiation in the NAIP (and BPIP as well), and the end of magmatism for volcanic products found across NAIP.
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2. SCIENTIFIC FRAMEWORK
This chapter serves as a literature study where the current state-of-the-art knowledge beyond the area of studied locations is defined. To understand the role of Lundy and Fur it is important to define the setting of the NAIP and BPIP and how their activity potentially caused climate change: the Paleocene- Eocene Thermal Maximum. This chapter will introduce the timing, duration, and effects of PETM had on the environment, and present opposing theories of carbon sources that caused the ‘greenhouse’
effect.
The Paleocene- Eocene period: paleogeography and climate
Part of the Cenozoic Era (from Greek kainos = new and zoon = animal), Paleogene Period (also spelled Palaeogene Period, from Greek palaios = old and genes = born or clan) represents oldest of the three stratigraphic divisions of the Cenozoic Era spanning the interval between 66 Ma and 23 Ma (Vandenberghe et al. 2012). The Paleogene period is classified into 3 epochs, or 9 stages: Paleocene (66.0–56.0 Ma), Eocene (56.0–33.9 Ma) and Oligocene (33.9–23.0 Ma) (Figure 2.1). The Paleocene began with the Cretaceous-Paleogene (K-Pg) extinction event that caused the extinction of non-avian dinosaurs among other taxa. Throughout the Paleocene the climate changed considerably: From temperate climate to the prolonged global warming. The high latitude and polar regions were ice-free with rich assemblage of plant and animal population, while the latitudes of ~30° where roughly tropical- with high temperatures and high precipitation (Zachos et al. 2001; Gingerich 2006). This is in literature termed as hothouse climate: periods in the geological history marked by great global warming event as a response of the Earth system to rapid carbon injection into the atmosphere and oceans. Although each hothouse climate, or global warming event, is unique, they share a lot of common attributes:
a) Rapid warming with abrupt onset duration of 100 – 100 000 years;
b) A total duration of between than 0.1 and 2 Myr (with Permian-Triassic being an outlier);
c) Negative carbon isotope excursion (CIE) of usually less than 4 ‰ (Figure 2.1) (again P-T having larger CIE);
d) Reduction in oceanic oxygen leading to anoxia and/or euxinia;
e) Increase of atmospheric CO2 and ocean acidification;
f) Augmentation of hydrological cycles ,with drier regions getting drier and humid regions getting more humid resulting in increased continental erosion/ weathering rates as well;
g) Flora and faunal response to the abrupt climate change (Foster et al. 2018).
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Figure 2.1: Chronostratigraphic chart of Paleogene period (International Commission on Stratigraphy, 2021) and isotope curve graph. The curve graphs of for both δ 13C and 18O (‰) shows position of three hyperthermals (PETM, ETM2, ETM3) as a pronounced negative spike on the graph. The data is acquired analysing the benthic foraminifera from the ocean cores from Litter et al. 2014.
The arrangement of continents in this period was already similar to recent times (Figure 2.2). Globally, India had separated from Antarctica, opening the West Indian Ocean. On the other side of the globe, a seafloor spreading event between Greenland and Europe began. The start of continental rifting was associated with igneous activity that formed the North Atlantic Igneous Province (Torsvik et al. 2002).
The Eocene period is marked by Eocene Climactic Optimum (ECO), where effects of thermal maximum started to gradually subside, turning a ‘hothouse’ climate to ‘icehouse’ at the end of Eocene (review in Westerhold et al. 2020) (Figure 2.1). Although during this period the climate started to cool down to form the Antarctic icecaps in the earliest Oligocene (~35 Ma), the climate did experience thermal maximums (hyperthermals: ETM2, ETM3) (Figure 3), probably triggered by the orbital eccentricity. The eccentricity regulates the distance to the sun and variations in insulation (Zachos et al. 2010).
In this period India has collided with Asia forming the Himalayan Mountain range. Continental rifiting between Greenland and Scandinavia opened the North Atlantic Ocean (Müller et al. 2016) (Figure 2.2).
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Figure 2.2: Paleogeographic reconstruction of continents at different stages: Early Paleocene (65 Ma), Paleocene- Eocene (56Ma) and late Eocene (34 Ma). Since the Early Paleocene, Greenland drifted away from Eurasia, forming the North Atlantic Igneous Province. The global plate motion model is after Müller et al. 2016.
Paleocene- Eocene Thermal Maximum (PETM): Timing, duration, causes &
effects
As stated in the previous chapter, the geological history has experienced cycles of ‘hothouse’ to
‘icehouse’ climate. Superimposed on this climate cycling are a number of geologically abrupt events:
the hyperthermals. The most recent and well recorded is the hyperthermal at Paleocene-Eocene boundary.
The boundary is defined by the onset of a pronounced negative carbon isotope excursion (CIE) (Figure 2.1), which is an expression of the PETM, a great global warming event (overview in Storey et al., 2007a). The timing and the trigger of the PETM is still heavily debated (Zeebe & Lourens, 2019;
Frieling et al. 2019).
It was around 55.93 Ma that the event started (onset of PETM), with main PETM body (or PETM CIE) lasting for 120-200 kyr (~55.8 Ma) (Westerhold et al. 2018). The PETM CIE is characterised by ~1-5 kyr rapid onset, a stable period, ~100kyr, and a progressive recovery period to pre-PETM pCO2 values and temperatures (McInerney & Wing, 2011). Global average temperatures increased by 5-8 °C
(McInerney & Wing, 2011). As the most of hyperthermal events coincided with astronomical processes (e.g. Milankovic cycle), the PETM on the other hand is considered anomalous in its magnitude and duration (Zachos et al. 2010).
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For that reason, several origins have been advocated: from a bolide impact (Kent et al., 2003), volcanic/magmatic degassing (Gutjahr et al., 2017) to the thermogenic degassing of organic rich sedimentary basins initiated by contact metamorphism around sills (Svensen et al., 2004; Aarnes et al.
2010) (Figure 2.3). The last two are connected to the emplacement of the NAIP.
Recent research is debating on the effects of both volcanic/magmatic degassing and the methane clathrate (also called methane hydrate: CH4·5.75H2O) dissociation as the source of the massive and rapid input of isotopically depleted carbon (Dickens, 2011; Gutjahr et al. 2017; Frieling et al. 2019).
Whatever may be the case, the PETM coincides with the basaltic flood magmatism (Storey et al.
2007b; Jones et al. 2019) and the emplacement of magmatic sills (55.6 ± 0.3 and 56.3 ±0.4 Ma: Svensen et al. 2010), indicating that the NAIP may have had a role in triggering the PETM (Svensen et al. 2004;
Storey et al. 2007a; Gutjahr et al. 2017; Frieling et al. 2019).
Figure 2.3: Different emplacement methods of the NAIP and possible causes of climate change in the Paleogene. A) Shallow marine environment causes violent eruptions that leads to wide transport of tephra and dispersal of toxic gasses in the atmosphere. When the ash cloud reaches stratosphere (~10km) where wind current is at much higher speed (60 m/s or 200 km/h), it can be transported to large distances; B) Effusive flood basalt volcanism can also transport tephra and large amount of gasses over extended periods of time and on large distances, depending on the eruption mechanisms; a source of the massive and rapid input of isotopically depleted carbon C) Hydrothermal vent complexes that most notably releases methane hydrate: CH4·5.75H2O. Adopted from Jones et al. 2016; 2019.
9 The effects of extensive carbon release include:
In the ocean:
a) deep-ocean acidification, causing the lysocline (depth at which carbonate starts to dissolve) to shallow, dissolving deep water carbonates;
b) ocean anoxia, result of absent bioturbation and water mixing (Zachos et al. 2005);
c) change of ocean circulation patterns, (Nunes & Norris, 2006);
d) sea level rise, caused by the thermal expansion of seawater and the melting of ice caps (Slujis et al. 2006);
e) extinction of 30-50% of all benthic foraminifera species (Thomas, 1990).
On land:
a) explosion of mammalian diversification and migration northward (Gingerich et al. 2003);
b) increase of extreme weather events, generally very humid weather conditions (Gingerich et al.
2003). The humid climate increased the erosion rate and sediment transport, meaning that sediments from the PETM are enriched with kaolinite (Stokke et al. 2020c).
Following the main CIE body, carbon isotope values and temperatures returned to near pre-PETM pCO2 values over ~83 kyr (Murphy et al. 2010). There are three main mechanisms for the drawdown of carbon: a) increased marine export production (Bridgestock et al. 2019), b) increased carbon and phosphorus burial (Komar & Zeebe, 2017; Longman et al. 2021) and c) accelerated silicate weathering (Bowen & Zachos, 2010), the last being the primary mechanism on the long scale. This recovery after the PETM-CIE is rapid (~30 ka; Bowen & Zachos, 2010) compared to other hyperthermals, suggesting that there were several mechanisms active at the same time (Komar & Zeebe 2017; Bridgestock et al.
2019).
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The North Atlantic Igneous Province and British Paleogene Igneous Province:
History, composition, and ages
The North Atlantic Igneous Province (NAIP) is the largest and most well-studied LIP in the Cenozoic.
LIPs are defined as magmatic provinces with areal extents of >0.1 Mkm2, igneous volumes >0.1 Mkm3, with maximum activity of 50 Myr that are spanned through short duration (1- 5 Myr) pulses. During that time, a large proportion (>75%) of igneous volume was emplaced and are of specific intraplate tectonic setting and/or geochemical affinities (Bryan & Ernst, 2008). LIP provinces include continental flood basalts and associated intrusive rocks, volcanic passive margins, oceanic plateaus, submarine ridges, seamounts, and ocean basin flood basalts. Unlike the magmatism that is associated with plate tectonics that creates new crust exclusively in the ocean basins or at ocean margins, LIPs can form independent of plate setting: on continents, in the oceans, and along margins between the two, and either wholly within plates or at plate boundaries (Coffin & Eldholm, 1994; 2005; Bryan & Ernst, 2008; Ernst 2014). The NAIP is a province in the Northeast Atlantic, recently centred around the Icelandic plume (Figure 2.4). The NAIP has an estimated total volume of around 6.6 Mkm3 and it consists of extrusive rocks, intrusions, and volcanic centres (see Saunders et al. 1997; Storey et al.
2007b; Meyer et al. 2007; Horni et al., 2007). Volcanic rocks, mainly flood basalts, and associated intrusions and dykes extend over a vast area that includes east and west Greenland, the Vøring Plateau off Norway, the Rockall Bank, and the British Paleogene Igneous Province (Figure 2.4).
A compilation of geochronological data (Wilkinson et al. 2017) suggests that NAIP magmatism occurred in two main phases: the pre-rift, phase one in the early Paleocene (ca. 62- 58 Ma), separated by the Eocene syn-rift phase (ca. 56 Ma), and the post-rift, phase two (ca. 56-53 Ma). Magmatism is thought to be a result of initial arrival of plume-head material, the initial plate breakup and opening of the North Atlantic Ocean. The phase one (the pre-rift phase) activity in the NAIP is made of tholeiitic basalt magmatism, that include emplacement of continental flood basalts (e.g. Major lava fields of Skye and Mull in Scotland, and the Antrim plateau in Northern Ireland).
The phase two (post-rift) is the most voluminous phase accompanied by polygenetic volcanism.
Geochemical analysis (overview in Meyer et al. 2007) has shown that the voluminous phase (c. 56–54 Ma) showed compositions that overlap with those of asthenospheric melts (as represented by recent Icelandic basalts and North Atlantic mid-ocean-ridge basalt (MORB). This voluminous phase is characterized by an extensive volcanism which is recorded in the tephra layers found along the North Atlantic Margins, North Sea and over Northern Europe (Larsen et al. 2003; Stokke et al. 2020b).
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Figure 2.4: Paleogeographic reconstruction of the North Atlantic Igneous Province at ~56 Ma. The small map in upper left corner shows known Large Igneous Provinces, with marked position of NAIP. LIP coloured in red are continental flood basalt provinces/volcanic rifted margins; yellow are silicic LIP; Dark blue are oceanic plateaux/ocean basin flood basalt provinces. The orange box represents the British Paleogene Igneous Province. Yellow circles are position of studied locations. Figure adapted from Jones et al. 2019.
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The British Paleogene Igneous Province (BPIP) is a small, off-axis part of NAIP. The major units are comprised of flood basalts located in Ireland, Scotland, and offshore. Significant volumes of silica- rich magmas formed major intrusive complexes and granite intrusions within the BPIP. These include Lundy Island, Ardnamurchan, Rum, Eigg, Skye, Arran, Mull, Antrim, and Canna (Figure 2.5).
Characteristics of the dyke intrusion in the BPIP is their NNW-SSE to N-S trending orientation, a result of Alpine deformation in Paleocene (Cooper et al. 2012; Anderson et al. 2013; Anderson et al.
2018). Following the field evidence and theoretical models, before the opening of the eastern North Atlantic (56-53.5 Ma), the entire region was uplifted, all before the igneous activity (~63 Ma) (e.g.
Baffin Island, West and SE Greenland and the British Isles). Before the continental breakup, subsidence is expected, but in the case of the European and East Greenland shelves, evidence suggest that this did not occur (Meyer et al. 2007). The missing subsidence in the shelf areas (East Greenland and offshore UK) is, supported by the geodynamic models (Meyer et al. 2007), a result of magmatic underplating (White & McKenzie, 1989; Kent & Fitton, 2000).
The lithosphere in the British Paleogene Igneous Province was preserved close to the Paleogene magmatism (Kent & Fitton, 2000; Meyer et al. 2007). The stretching of continental lithosphere continued after c. 63 Ma and was later accompanied by the eruption of basaltic lavas across a large area stretching from Rockall and St Kilda, west of the Outer Hebrides, east to Skye, Mull, Rum, Muck, Eigg and Arran, south to Ardamurchan and Ailsa Craig to the Mourne Mountains, Slieve Gullion and Carlingford Lough in NE Ireland (Ritchie & Hitchen, 1996; Brown et al. 2009). Previous work on geochemistry and Pb-Nd-Sr isotopic studies of mafic lavas showed four magma types: M1-M4, that is the transition between tholeiitic to mildly alkalic basaltic rocks (review in Kent & Fitton, 2000). They proposed that two mantle sources were present during the Paleogene: 'Icelandic plume' and a normal (N)-type mid-ocean ridge basalt (MORB) which is a distal part of ancestral plume. Although majority of BPIP comprises of these types of rocks, an important, though volumetrically minor component is granite intrusions scattered in the BPIP. An important part of research of the BPIP is dating the different dyke sets which would provide temporal constraints for pulses of extension and rifting, magma initiation, and the end of magmatism for volcanic products found across NAIP. To compile this information it is crucial to use most precise dating techniques and regularly update the dataset.
This thesis plays the role in the said update, as it will present new and improved age data compared to those from Charles et al. (2017). The most recent review of the geochronological data set can be found in Wilkinson et al. 2017 (Figure 2.5).
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Figure 2.5: The compiled age map of British Paleogene Igneous Province and its main igneous intrusions.
The complied age data can be found in Wilkinson et al., 2017. To increase the precision and better constrain the resolution of magmatic/volcanic activity it is important to use currently most precise dating methods: U-Pb zircon on felsic intrusions and Ar/Ar on basic intrusions. The age of intrusive complexes are from: Rockall seamounts- O’Connor et al. 2000; St. Kilda- Brook, 1984; Lewis- Faithfull et al. 2012; Skye & Rum- Hamilton et al. 1998; Muck- Chambers et al. 2005; Mull- Chambers & Pringle, 2001; Arran- Musset et al. 1987; Ailsa Craig- Harrison et al. 1987; Antrim- Ganerød et al. 2011; Slieve Gullion, Mourne and Carlingford- Gamble et al. 1999; Lundy- Charles et al. 2017.
Asterisk is on Lundy age to indicate that this study is going to provide with new age.
Impact of NAIP: can volcanism cause climate change?
Most LIP activity is marked by the intensive volcanism and subsequent volcanic gas release (CO2, SO2) which acts as a catalyst for environmental crises (review in Bond & Wignall, 2014; Robock, 2000). Although, there is no direct correlation between total volume of lava and released gas (Jones et al. 2016, Jones et al. 2019) with the mass extinctions, the temporal link between environmental crisis and LIP emplacement is well known. This is well researched with the Phanerozoic ‘Big 5’ mass extinction events: Late Devonian (~372 Ma), Middle Permian (~260 Ma), end- Permian (~252 Ma), end-Triassic (~201 Ma) and Early Jurassic (~183 Ma) (Bond & Wignall, 2014).
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In the case of PETM, the hyperthermal event is synchronous with the voluminous (flood basalt) phase of magmatism (phase two) (Meyer et al. 2007; Storey et al. 2007b; Jones et al. 2019; Stokke et al.
2020b) and emplacement of magmatic sills (~55.6 Ma; Svensen et al. 2010). As mentioned, two proposed mechanisms that can increase the temperature and overall cause perturbations in the geosphere from LIP activity: with volcanic/magmatic degassing through flood basalt eruptions (e.g.
Gutjahr et al. 2017) and through degassing generated by contact metamorphism in the sills found in the sedimentary basins (Svensen et al. 2004, Aarnes et al. 2010).
The main emitted volcanic gases from basaltic eruptions are H2O, CO2, SO2, with HF, HCl and poisonous gases that with large-magnitude eruptions cause great climate disturbances. Of these gases only CO2 causes significant warming over the longer period (Jones et al. 2016). Carbon, a greenhouse gas, is present in large quantities in surface reservoirs where the cycle depends on the fluxes between theses reservoirs. For this reason, it is believed if the large-scale degassing (i.e. mantle degassing or thermogenic degassing) happens over a short time, it could be a potential trigger for the PETM warming event (Gutjahr et al. 2017).
Contrarily, SO2 can cause localized short-term warming but its major effect is cooling as this gas forms sunlight-blocking aerosols. Famous example of this effect is recorded in the recent human history with the eruption of Mount Pinatubo in 1991 (Hanses et al. 1992). Callegaro et al. 2014 suggested, by measuring the concentrations of sulphur in melt inclusions, that previous LIP events that were associated with mass extinctions were sulphur rich. This, off course, is hard to absolutely determine as the residence time of sulphur is short, compared to carbon (Jones et al. 2016).
On the other hand H2O and HCl are known to condense on ash particles to from acidic rains that can affect flora and fauna. Additionally, HCl destroys protective ozone layer (Bond & Wignall, 2014).
Recent example of ecological failure on land is the 1783-1784 eruption of Laki volcano on Iceland (Thordarson & Self, 2003). The Laki event lasted just 8 months, and yet caused crop failures and hundreds of thousands of deaths across Europe and into North Africa. This minor geological event caused widespread environmental damage, and the volcanism was around 450,000 times smaller than that of the NAIP. Although the residence time of HCl is short, in combination with greenhouse emission, it can trigger mass extinction, with one good example from the end- Permian mass extinction found in rock record of Siberian Traps (Black et al. 2013).
The current challenge is to better understand the variable environmental effects of LIP eruptions and identify why these mechanism(s) sometimes cause catastrophic extinctions.
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3. STUDY LOCATIONS
This chapter presents the overview of the research done on the two of the studied locations:
Lundy Island and Fur Island. The chapter will give overview of the petrological and geochemical characteristics of both the granite, the accompanied dykes and the fur ashes (especially ash layer -33) with goal to define the similarities and differences of materials from these two locations.
Geology of Lundy Island
Figure 3.1: A geological map of Lundy. Most of the island is made of granite (orange), while the southern tip is a part of Devonian slate series (grey-not researched in this study).
The entire island is cut by various mafic and felsic dykes that post-date the main granite body. Data on magnetic dykes is from Roberts & Smith, 1994. Map is adapted from Charles et al., 2017.
Lundy (ca. 4.5 km2) is situated in the Bristol Channel, 20 km NNW of the north Devon coast (Figure 2.5). The geological map of the Lundy shows that the island is relatively homogeneous (Figure 3.1). It is composed (>90%) of two mica±
garnet ± tourmaline granite with peraluminous signature s-type granite signature. The SE part of the island are made of the low-grade metasedimentary rocks, mainly thick, folded, and well-jointed arenaceous slates of Devonian age.
Both are cut by suite of basalt-dolerite, trachyte, and rare rhyolite dykes (Figure 3.1) (Dollar, 1941; Thorpe & Tindle, 1992; Thorpe et al. 1990;
Charles et al. 2017).
Dollar, (1941), correlated the slate series with the upper Devonian Morte Slates of north Devon.
The boundary between the granite and the slates follows NNE-SSW trend and is partially fault controlled. The fault plane displays two sets of striation, indicating strike-slip movement (Charles et al. 2017).
Initial research on granites (Dollar, 1941) described two main types: medium- to coarse-grained megacrystic two mica granites (G1 and G2). Charles et al. (2017), showed that it is very hard to distinguish those types of granites in the field. This classification is better observed microscopically.
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The Lundy dykes are predominantly olivine basalts-dolerites with subordinate peralkaline trachytes and rhyolites containing K-feldspar + quartz + amphibole ± clinopyroxene. Dykes display a variation in grain size and texture, from porphyritic to aphyric. The granites are cross cut by pegmatites and microgranite sheets. Small xenoliths of the slates appear in the granites at the granite- metasediments boundary in the SE of the Island. Edmonds et al. (1979), suggested that in addition to the dextral strike- slip faulted boundary, there might be an extensional contact, but that junction is cut by heavily brecciated basic dyke. The faulted contact between granite and the slate series is consistent with the lack of thermal metamorphism in the metasedimentary rocks and absence of granite veining. There are over 200 counted dykes (Dollar, 1941) with average thickness of 1m and varying composition from intermediate to basic (90% of the dyke swarm is basic), cross cutting the granite and Lundy Slate Series. Most dykes are doleritic in composition, but some felsic dykes appear, dominated by quartz trachytes. The dykes post-date the solidification of the granite (Charles et al. 2017).
Modern geochemical studies have investigated the roles of mantle and crustal melting, differentiation, and contamination in generating the Lundy granite (Stone, 1990; Thorpe et al. 1990) and basalt- trachyte-rhyolite dykes (Thorpe & Tindle, 1992). Yb, Nb and Rb trace element data for granites indicate an affinity with ‘syn-collisional’ or ‘within-plate’ granites (Stone, 1990; Thorpe et al. 1990).
87Sr/86Sr ratio of ca. 0.715 and εNd values of -0.9 to -1.9 show evolved Sr and Nd signature, plotting between positive εNd values of mantle and highly evolved granites of BPIP (e.g., Cornubian granites:
εNd = -8 to -9; Stone, 1990), suggesting a crustal source to the Lundy granite. The average 87Sr/86Sr ratios for dolerite-trachyte-rhyolite dykes is 0.705 and εNd values range from 5.9 to 9.0 (for dolerites) to 7.6 (trachyte) and 6.0 to 7.9 (for rhyolites) are suggesting the later influx of mantle material after the emplacement of granite (Thorpe & Tindle, 1992). Thorpe et al. (1990), proposed that Lundy granite was derived from mixing of parental magma consisting of crustal component and a mantle derived component derived from a basaltic magma. Following emplacement of the granite, the basaltic magma chamber remained active and was sampled during further fractional crystallization to create dolerites- trachyte-rhyolite magmas emplaced as dykes into the cooler, jointed granite (Thorpe & Tindle, 1992).
The appearance of the Lundy granite has been connected to a Paleocene NNW-SSE dextral fault system which are result of N-S Alpine shortening (i.e. from convergence of Africa and Europe, Cooper et al. 2012; Anderson et al. 2013; Anderson et al. 2018), possible localized rifting (Charles et al. 2017) and distal magmatic processes of the ancestral Iceland plume (White & McKenzie, 1989; Kent &
Fitton, 2000). Regionally, this NNW-SSE striking faults are evident in the Northern Ireland and the Irish Sea, where the Codling Fault is displaying 9 km of dextral displacement (Anderson et al. 2013).
Displacement is further transferred onto the Sticklepath-Lustleight Fault (Dart et al. 1995), where Lundy is geographically positioned (Figure 3.2).
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Several dating isochrones were used for the Lundy granite: Rb-Sr whole rock isochrones gave 58.7±
1.6 Ma (Thorpe et al. 1990), combined with earlier whole rock and mineral radiometric dating of granites and dolerites (Edmonds et al. 1979). The most recent dating using U-Pb zircon ages resulted in the more precise ages of 59.8 ± 0.4 – 58.4 ± 0.4 Ma (Charles et al. 2017). This indicates that that Lundy Igneous Complex is the most southerly expression of the British Paleogene Igneous Province and the wider North Atlantic Igneous Province (Wilkinson et al. 2017).
Figure 3.2: Distribution of dyke swarms and major faults in the greater area of BPIP. Black dashed lines are dyke swarms and red lines represent position of major faults. Lundy is positioned adjacent to the Sticklepath Fault (SF), a Paleocene NNW-SSE dextral fault system that transitions into Codling Fault (CF). Compiled data is from Cooper et al. 2012 and British Geological Survey: Geology of Britain viewer.
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Geology of Fur Island
Figure 3.3: The elevation map and photo of the locality of Stolleklint, with PETM onset/recovery and ash 33.
The ash -33 is a boundary layer between Stolleklint Clay and Diatomite of Fur Formation (see next figure). Figures from Stokke et al. 2020a.
Fur is a small island (22km2) located in the NW Denmark’s Limfjorden area. An entire section of Paleocene- Eocene stratigraphy is exposed and easily accessible in the coastal area (Figure 3.3). The Fur Ash layers are found in the Fur diatomite, which lies within the Ølst Formation, described as a sequence of clays interbedded with volcanic ash (Heilman-Clausen et al. 1985; Pedersen & Surlyk, 1983). The Fur Formation lies on the Stolleklint Clay where the beginning of PETM is identified (Figure 11) (4.5 ‰ CIE, Jones et al. 2019). The clays are described as sandy, silty and non-calcareous of dark-grey colour (Heilman-Clausen et al. 1985; Pedersen & Surlyk, 1983). The Stolleklint Clay is an important stratigraphic layer as it correlates to the Sele Formation of offshore North Sea Basin (Figure 2). It was interpreted that the Stolleklint Clay and Fur formation diatomite are deposited during the progressive marine transgression at the Paleocene-Eocene boundary between bathyal and outer neritic (Rasmussen et al. 2008, Schoon et al. 2015). The reason for the sea level change in the basin is caused by the activity of Icelandic plume and the start of sea-floor spreading (Schoon et al. 2015).
Volcanic ash layers found in the Fur Formation (ca. 200 ash layers) are divided in the ‘negative’ and
‘positive’ series, based on their petrological composition (Figure 3.4) (Bøggild, 1918). Petrological analysis and early major-element analyses (e.g. Pedersen & Surlyk, 1983; Morton & Evans, 1988) recognise four stages of volcanic activity (Pedersen et al. 1975; Larsen et al. 2003).
Stage 1 correlate with the ash layers from -39 to -22, varying composition from rhyolitic to basaltic.
The sources of these ash layers are thought to be from volcanic centres within BPIP (Larsen et al.
2003).
19 Figure 3.4: The detail of stratigraphy at
Stolleklint. Gray lines indicate the position of each ash layer and their thickness at Fur Formation (Bøglide, 1918). Orange box shows extent of PETM carbon isotope excursion (Stokke et al. 2020a). 1Charles et al. 2011, if timing of Svalbard’s and Fur CIEs are coeval; 2Storey et al. 2007a;
3Westerhold et al. 2009; 4Chambers et al.
2003. Figure from Stokke et al. 2020b.
The ash layer -33, the focus of this study, a peraluminous rhyolite shows a strikingly similar geochemical signature with the S-type granite found on Lundy Island (Thorpe &
Tindle, 1992). This layer is glass-rich, and the glass is fresh (Pedersen et al., 1975). In the context of PETM layer -33 marks the end of the CIE body (Figure 3.4) (Jones et al.
2019). Recent research on the role of ash -33 in the end of CIE, showed that carbonate cement found within this layer, may play an important role in mitigating the excess carbon from atmosphere (Longman et al. 2021). Radiometric dating of an ash layer within the recovery of the PETM CIE in Svalbard yielded age of ~55.785 Ma (Charles et al. 2011), which would be close to the stratigraphic position of ash -33 if the CIE’s in Fur and Svalbard are coeval.
Stage 2 corresponds to the ash layers -21 to -15, variable as well (phonolites, nephilinites, trachytes and rhyolites). The characteristic property of the negative series is their extreme compositions and the high level of alteration. Interesting layers from this stage are the peralkaline nephelinitic -19, recognized as a blue ash layer in the field and the diagenetically altered trachyte, orange coloured -17 ash layer (Larsen et al. 2003). In the PETM event, layer -21a is important as it represents the final end of recovery (Figure 3.4) (Jones et al. 2019).
Stage 3 is represented by the three distinctive black alkali basalt ash layers -13, -12 and -11. They are the least altered in the negative ash series, and their major-element diagrams form a continuation of the trends observed in the positive series (Larsen et al. 2003). These ash layers, if compared to other rocks in the North Atlantic, show similarities to the alkali basalts from both East Greenland and Iceland (Larsen et al. 2003).
Stage 4 ash layers belong the positive ash series, which are relatively uniform in their geochemical signature. Most ash layers are tholeiitic ferrobasalts, with exception of layers +13 and +19 being sub- alkaline rhyolites. Ash +19 contains fresh glass that is not heavily altered (Pedersen et al. 1975).
Tholeiitic basalts correlate to the basalts currently forming in Iceland while rhyolites might either correlate to the central volcanic complexes in Iceland or dykes found in the BPIP, including Lundy as a potential source (Larsen et al. 2003).
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4. THEORETICAL BACKGROUND AND INSTRUMENTATION
This chapter will give an overview on theory behind the methods and result chapter. This chapter will start with principles of geochemistry: application of trace and rare earth elements in petrology. It will continue onto principles of radioactivity and description of the U-Pb method for which the resulting ages are based on. The chapter will end with short background on the how instrumentation works.
There are many textbooks and articles on the (isotope) geochemistry, therefore all the information, if not specifically stated, is sourced from Faure & Mensing, (2005) and Gill, (2011).
Geochemistry: trace elements as ‘fingerprints’ in geology
Generally, in geochemistry, trace elements are defined as elements whose concentration is less than 1000ppm (parts per million) or 0.1% of a rock’s composition. This definition is considered too general, as certain rocks can contain more than 1000ppm of a trace element, and still be considered a trace element. Modern geochemical analysis usually includes oxides (Al, Ca, Fe2+, Fe3+, K, Mg, Mn, Na, P, Si, Ti), REEs (Rare Earth Elements: Lanthanides and Actinides plus Y), and series of other trace elements, whose concentrations vary more widely than the major elements. Trace elements can further be classified based on their ionic radii, cation charges and partition coefficient (Kia) (Figure 4.1). This separates incompatible trace elements from compatible ones. Incompatible elements favour the melt over the coexisting mineral crystals. Contrarily, compatible elements favour crystallized structures over the melt. Furthermore, incompatible elements can be divided into Large Ion Lithophile Elements (LILE) and High Field Strength Elements (HFSE) (Figure 4.1). The LIL elements have too large of ionic radius to be accommodated in most rock forming minerals. These elements are, due to their property, often mobile during weathering or alteration. The HFS elements have a high charge to radius ratio, making them unstable in ionic silicate crystals. They are less mobile and less prone to dissolution (e.g. weathering or alteration). Compatibility also depends on the composition of the melt. For example, elements that are incompatible (Kia < 1) in the mafic (silica SiO2-poor) melts will have coefficient =/>1 in silicic (SiO2-rich). This is best illustrated with solubility of Zr in zircon (ZrSiO4):
in basaltic melts zircon will usually not crystalize while in granitic melts it will readily crystallize as a phase. Trace elements are, therefore, useful geochemical discriminants as they can provide useful information on the origin of the rocks, their evolution, and their tectonic setting.
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Figure 4.1: A diagram of compatible vs. incompatible elements. Diagram is plotted as ionic radius vs. cation charge.
Based on this diagram elements are placed into two groups: incompatible and compatible plus the boundary where elements can fall into either group. Italics: LILE; bold: HFSE; Large: major elements. The box represents 15 of REEs together with europium (Eu), who is unique in behaviour. b) Cartoon depicting compatible and c) incompatible behaviour in a mineral. Kia is the partition coefficient, indicating the affinity towards the crystal structure. Sourced from Gill, (2011).
Isotopes, radioactivity, and decay
The nucleus of an atom is made of nucleons of which there are two kinds: protons and neutrons. The total number of nucleons in the atomic nucleus is called the mass number (A). The number of protons is called the atomic number (Z). The number of neutrons in the atomic nucleus ranges in value for any given element, corresponding to different isotopes of said element. For example, 816O is an isotope of oxygen with 16 nucleons of which 8 are protons (and N = A-Z = 16-8 = 8 are neutrons). When another neutron is added to the atomic nucleus of oxygen, the nucleus becomes unstable and undergoes radioactive decay.
Radioactivity is a process where an unstable atomic nucleus spontaneously transforms into other nuclei by giving off particles and energy in the form of radiation. This phenomenon can be well observed in the chart of nuclides. The chart of nuclides is made of proton/atomic number (Z) on the y-axis and neutron number (N) on the x-axis (International Atomic Energy Agency, IAEA) (Figure 4.2). The unstable nuclides (or radionuclides), such as 82209Pb or 819O may survive for time periods of femtoseconds to billions of years depending on the degree of instability, which generally scales with the ‘distance’ from the curve of stable nuclides. Radionuclides eventually disintegrate to a stable form by means of several different mechanisms: α-decay, β-decay, electron capture or/and nuclear fission.
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Figure 4.2: The distribution of natural stable isotopes in the neutron-proton diagram. Also called chart of nuclides. Natural (or artificial) radioactive isotopes lie between the stable isotopes. After N=20 the line of stable nuclides shifts from the diagonal where the number of proton and neutrons are the same. For N>20, number of neutrons are greater than the number of protons, which makes heavy isotopes decay easily. The orange line (Z=N) is termed the valley of stability as it corresponds to the minimal energy to keep isotopes stable. Figure from White, (2015). The data on each isotope can be found at official pages of IAEA using their Live Chart of Nuclides.
Principles of U-Pb geochronology
A characteristic property of radioactive decay is its absolute independence of external physical and chemical effects. In other words, it is not affected by changes in pressure, temperature, or the molecular bonds connecting a radioactive nuclide to neighbouring atoms. This means that the rate at which a radioactive parent decays to a radiogenic daughter per unit time, i.e., dN∕dt only depends on N, the number of parent atoms present. The time it takes for a radioactive isotope to be reduced by half is called half-life, defined by 𝑇1/2= 𝑁0/𝑁
𝜆 [𝑦−1], where N is the number of the remaining radioactive nuclei, N0 is the initial number of radioactive nuclei and the λ, decay constant of radioactive nuclei.
The U-(Th)-Pb geochronology is in some ways special, as it decays through several mechanisms from multiple radioactive parent isotopes (238U, 235U, 232Th) to different stable isotopes of Pb (206Pb, 207Pb and 208Pb), each with its own half-lives and decay-chains (Figure 4.3).
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Figure 4.3: The chart of nuclides showing the decay series of 232Th, 235U and 238U. These isotopes form the basis of the U-Th-Pb, U-Th-He and U-Th-series methods. The number inside boxes represent half-life of each isotope while Greek alphabet symbols represent mode of decay. From 232Th we get 208Pb; 235U we get 207Pb and from 238U we get 206Pb.
Figure from White, (2015).
Each of these three decay series is unique, i.e. no isotope occurs in more than one series (Figure 4.3).
Furthermore, the half-life of the parent isotope is much longer than any of the intermediary daughter isotopes, thus fulfilling the requirements for secular equilibrium.
Natural Pb consists of four isotopes 204Pb, 206Pb, 207Pb and 208Pb. The ingrowth equations for the three radiogenic Pb isotopes are given by:
1.1)
With λ238 = 1.55125 ×10-10a-1 (t1∕2 = 4.468 Gyr), λ235 = 9.8485 ×10-10a-1 (t1∕2 = 703.8 Myr), and λ232 = 0.495 ×10-10a-1 (t1∕2 = 14.05 Gyr) (Steiger & Jäger, 1977).
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Some igneous minerals (notably zircon) conveniently incorporate trace amounts (typically hundreds to thousands of ppm) of U and virtually no Pb upon crystallization. This is due to similar properties of Zr4+ and U4+ (similar ionic radius and cation charge), as these elements can replace each other while Pb2+ is excluded from the structure during crystallization because of its higher ionic radius and low charge (Figure 4.1). For zircon, the non-radiogenic Pb can be safely neglected so that we can assume that Pb ≈ Pb*. This assumption cannot be made for other minerals, young ages, and high precision geochronology. In those cases, the non-radiogenic component (aka ‘common Pb’) needs to be quantified, which is done by normalising to non-radiogenic 204Pb:
1.2)
Where ∘ stands for the common Pb component for isotope x.
Zircon based, U-Pb dating grants access to two separate geochronometers (206Pb/238U and 207Pb/235U) based on different isotopes of the same parent-daughter pair (i.e., U & Pb). Two independent ages can be determined, based on the decay of two radioactive U isotopes 235U and 238U, and accumulation of radiogenic 207Pb and 206Pb (Formula 1.3). Comparison of 206Pb/238U and 207Pb/235U ages provides a test for closed system behaviour. If the two independently determined ages agree within uncertainty, the date is called concordant, if they do not agree the date is called discordant. The initial Pb composition can either be determined by analysing the Pb composition of a U-poor mineral (e.g., galena or feldspar) or by applying the isochron method to samples with different U and Th concentrations (Schoene et al.
2013).
From Equation 1.1, we find that:
1.3)
If 206Pb*/238U- and 207Pb*/235U-ratios which yield the same ages (t) are plotted against one another, they form a concordia curve (Figure 4.4).