UNIVERSITY OF HELSINKI
DEPARTMENT OF PHYSICS
REPORT SERIES IN GEOPHYSICS
No 77
BALTICA FROM THE SVECOFENNIAN UNTIL
THE EDIACARAN: A PALEOMAGNETIC STUDY
OF PRECAMBRIAN FORMATIONS OF FINLAND
Robert Klein
Cover images:
Top left: Geomagnetic field
Top right: Åland diabase dyke
Bottom left: Satakunta sandstone (Satu Mertanen)
Bottom right: Laurentia, Baltica, Amazonia and Siberia in Nuna configuration
HELSINKI 2016
Supervisors:
Prof. emer. Lauri J. Pesonen
Department of Physics
University of Helsinki
Helsinki, Finland
Dr. Satu Mertanen
Geological Survey of Finland
Espoo, Finland
Prof. Ilmo Kukkonen
Department of Physics
University of Helsinki
Helsinki, Finland
Pre-examiners:
Dr. Sergei Pisarevsky
School of Earth and Environment
University of Western Australia
Crawley, Australia
Prof. Phil McCausland
Department of Earth Sciences
Western University
London, Ontario
Canada
Opponent:
Custos:
Prof. David A.D. Evans
Department of Geology and Geophysics
Yale University
New Haven, USA
Prof. Ilmo Kukkonen
Department of Physics
University of Helsinki
Helsinki, Finland
Report Series in Geophysics No. 77
ISBN 978-951-51-1945-2 (printed version)
ISBN 978-951-51-1946-9 (pdf version)
ISSN 0355-8630
Helsinki 2016
Unigrafia
BALTICA FROM THE SVECOFENNIAN UNTIL THE
EDIACARAN: A PALEOMAGNETIC STUDY OF
PRECAMBRIAN FORMATIONS OF FINLAND
Robert Klein
ACADEMIC DISSERTATION IN GEOPHYSICS
To be presented, with the permission of the Faculty of Science of the University of Helsinki for
public criticism in the Auditorium E204 of Physicum, Gustaf Hällströminkatu 2a, on 11 March,
2016, at 12 pm.
Helsinki 2016
Table of contents
List of original articles
7
Abstract
8
Acknowledgments
11
1. Introduction
12
1.1. Supercontinents
12
1.1.1. Nuna
13
1.1.2. Rodinia
15
1.2. Paleomagnetism
17
1.2.1. The geocentric axial dipole
17
1.2.2. Magnetism in rocks
18
1.2.3. Paleomagnetic measurements
21
1.2.4. Apparent polar wander paths and paleogeographic reconstructions
22
1.2.5. Mafic dykes and sedimentary rocks in paleomagnetic studies
23
2. Aims of this study
25
3. Materials and methods
25
3.1. Sampling and sample preparation
25
3.2. Laboratory measurements.
26
3.2.1. Petrophysical measurements
26
3.2.2. Paleomagnetic and rock magnetic measurements
27
28
4. Review of articles
4.1. Article I
28
4.2. Article II
29
5
4.3. Article III
30
4.4. Article IV
31
4.5. Author’s contribution
32
33
5. Results and discussion
5.1. Paleomagnetic results obtained in this work
33
5.1.1. Diabase dykes
33
5.1.2. Sedimentary rocks
36
5.1.3. Secondary component
36
5.2. Apparent polar wander paths for Baltica (1.87 – 0.57 Ga)
38
5.3. Paleogeographic reconstructions (1.87 – 0.57 Ga)
41
6. Conclusions
50
7. References
52
6
List of original articles
This thesis is based on the following four articles, which are referred to in the text by their Roman
numerals:
I.
Klein, R., Pesonen, L.J., Salminen, J., Mertanen, S., 2014. Paleomagnetism of
Mesoproterozoic Satakunta sandstone, Western Finland. Precambrian Research 244, 156169.
II.
Salminen, J.M., Klein, R., Mertanen, S., Pesonen, L.J., Fröjdö, S., Mänttäri, I., Eklund, O.,
2015. Palaeomagnetism and U–Pb geochronology of c. 1570 Ma intrusives from Åland
archipelago, SW Finland – implications for Nuna. In Li, Z. X., Evans, D. A. D. and Murphy,
J. B. (eds) Supercontinent Cycles Through Earth History. Geological Society, London, Special
Publications 424, http://doi.org/10.1144/SP424.3.
III.
Klein, R., Salminen, J., Mertanen, S., 2015. Baltica during the Ediacaran and Cambrian: A
paleomagnetic study of Hailuoto sediments in Finland. Precambrian Research 267, 94-105.
IV.
Klein, R., Pesonen, L.J., Mänttäri, I., (submitted). A new late Paleoproterozoic pole for
Baltica: A paleomagnetic and geochronological study of the Keuruu diabase dykes, Central
Finland. Precambrian Research.
Article I and III are reprinted from Precambrian Research with the permission of Elsevier. Article II
is reprinted from Geological Society of London Special Publications with permission from the
Geological Society.
7
Abstract
This dissertation presents paleomagnetic results from Baltica for three significant time intervals
during the evolution of Precambrian supercontinents: i) 1.87 Ga – the onset of the supercontinent
Nuna; ii) 1.58–1.57 Ga – maximum compaction of Nuna; and iii) ca. 0.6–0.57 Ga – the break-up
between Baltica and Laurentia which marked the final break-up of the supercontinent Rodinia.
Paleogeographic reconstructions at 1.25 Ga and 1.05 Ga – which marks the transition from Nuna to
Rodinia - is briefly discussed based on existing paleomagnetic results; no new data were produced
for these two reconstructions.
New paleomagnetic results were obtained from samples from the 1.87 Ga Keuruu diabase dyke
swarm (Central Finland), 1.58-1.57 Ga Åland intrusives (Åland archipelago, SW Finland) and ca.
0.6-0.57 Ga Hailuoto sedimentary formation (Western Finland). In addition to these, new
paleomagnetic results were obtained from the Satakunta sandstone formation (SW Finland),
providing a new paleomagnetic age estimate of ca. 1.63 – 1.45 Ga for the sandstone deposition, older
than previously thought (ca. 1.4-1.3 Ga).
The 1.87 Ga Keuruu diabase dykes yield a dual polarity intermediate NNW downward or SSE upward
characteristic remanent magnetization (ChRM), positioning Baltica at low latitudes in a way that
would proceed into the 1.8-1.2 Ga NENA (North Europe – North America) configuration. The age
of the Keuruu dykes are defined by a U-Pb (zircon) age of 1870 ± 9 Ma and a Pb-Pb (zircon) age of
1869 ± 7 Ma. The primary nature of the Keuruu ChRM is supported by positive baked contact tests.
The Åland intrusives and Satakunta sandstone both yield dual polarity shallow (upward and
downward) NNE or SSW characteristic remanent magnetizations (ChRM), positioning Baltica on
equatorial paleolatitudes and supporting the NENA configuration between Baltica and Laurentia
during the Mesoproterozoic. Well determined U-Pb (zircon) ages narrows the occurrence of intrusion
to 1570 – 1580 Ma. The primary nature of the Åland intrusives ChRM is supported by positive baked
contact tests.
The Hailuoto sedimentary samples were obtained from a 272 m deep fully oriented drill core. The
depostional age of the Hailuoto sediments is poorly constrained at 570-600 Ma, based on fossil
records. The Hailuoto pole adds to the scattered Ediacaran paleomagnetic data of Baltica and indicates
large distances between other late Neoproterozoic and early Cambrian paleomagnetic poles, implying
rapid movement of Baltica from high latitudes at 615 Ma, over the polar region, to low latitudes at
550 Ma.
8
In all these cases a secondary component with shallow to intermediate downward inclinations and
northeasterly declinations were observed. This secondary component is commonly observed in the
Fennoscandian Shield and has previously been interpreted as a result of oxidizing fluids during the
Phanerozoic. In the case of Keuruu and Åland, meteorite impact (from Keurusselkä and Lumparn,
respectively) cannot be ruled out as a possible cause of the secondary component. This component is
mostly observed at low coercivities/blocking temperatures, but sometimes also at higher
coercivities/blocking temperatures, causing asymmetry in the dual polarity results of Keuruu diabase
dykes and Åland intrusives.
In summary, the work presented in this dissertation provides new paleomagnetic results for Baltica,
adding two new key poles to the Paleoproterozoic-Mesoproterozoic apparent polar wander path
(APWP) of Baltica. Furthermore, the results provide a better understanding of the movement of
Baltica during the enigmatic Ediacaran period. Paleogeographic reconstructions are presented at 1.87
Ga, 1.57 Ga, 1.25 Ga, and 1.05 Ga. The 1.87 Ga reconstruction do not support the NENA or SAMBA
configurations, but indicates the onset of Nuna. The 1.57 Ga and 1.25 Ga reconstructions indicate
similarities in the relative positions between Baltica, Laurentia, Amazonia, and Siberia during this
period. The transition between Nuna and Rodinia is marked by a rotation between Laurentia, and
Baltica and Amazonia (and possibly West Africa), forming the Grenville collisional belt.
9
10
Acknowledgments
Support for this work was provided by the Finnish Doctoral Program in Geosciences (formerly the
Finnish Graduate School in Geology), and the Vilho, Yrjö and Kalle Väisälä Foundation. Additional
funding for field work and travel was provided by the Satakunta Regional Fund of the Finnish
Cultural Foundation, the Oskar Öflund Foundation, and the University of Helsinki Chancellor’s travel
grant.
I would like to express my gratitude to my supervisors, Prof. emer. Lauri Pesonen, Dr. Satu Mertanen,
and Prof. Ilmo Kukkonen, for their support during my thesis work. I would especially like to thank
Prof. emer. Pesonen for giving me the opportunity to work at the Solid Earth Geophysics Laboratory.
I sincerely thank the pre-examiners of this thesis, Dr. Sergei Pisarevsky (University of Western
Australia), and Prof. Phil McCausland (Western University, Canada) for their constructive reviews,
which helped to improve this thesis significantly.
A special thanks to my co-authors, and colleagues (former and current), without whom this work
would not have been possible. I am especially indebted to Johanna Salminen and Tomáš Kohout for
their scientific and technical support.
Finally, I would like thank my family and friends for their moral support. My warmest gratitude goes
to my wife, Anu, for her support and patience, especially during the final stages of my thesis.
11
1. Introduction.
1.1. Supercontinents
Supercontinents are assemblies that contain all, or most of the Earth’s continental blocks (Rogers and
Santosh, 2004). The most recent supercontinent, Pangea, reached its maximum packing at ca. 250
Ma (e.g. Torsvik et. al., 2012), after which it started to rift apart, forming our present ocean basins.
The configuration of Pangea can be best determined by tracing the patterns of magnetic stripes in the
oceans that opened within it. Pangea consisted of Gondwana in the south, and Laurasia in the north.
Gondwana comprised the southern continents - South America, Africa, India, Madagascar, Australia,
and Antarctica - and had already formed at ca. 550 Ma (Meert and Van der Voo, 1997), and accreted
to Pangea as a single block. The concept of Gondwana originated when paleontologists noted that
Paleozoic and Early Mesozoic fossils from South America, Africa, India and Australia were all very
similar. Laurasia consisted of Laurentia (core of the North American continent) and Eurasia, however
a large part of it formed during the accretion of Pangea (Rogers and Santosh, 2004).
Several pre-Pangea supercontinents have been proposed. Apart from paleomagnetic data (explained
below), there is also geological evidence of the existence of Precambrian supercontinents. Orogenic
belts are thought by some to have developed during the formation of supercontinents, and are now
scattered over the present day continents. The Paleo- to Mesoproterozoic supercontinent Nuna (also
called Columbia; e.g. Zhao et al., 2004) was reconstructed largely on the basis of aligning the 2.1-1.8
Ga orogenic belts that occur on many continental blocks. Evidence for the late Meso- to
Neoproterozoic supercontinent, Rodinia, is partly based on the presence of 1.1-0.9 Ga orogenic belts
(Hoffman, 1991). Also, dyke swarms provide evidence of the existence of Precambrian
supercontinents where they occur on several continents that have rifted apart after the dyke intrustion
(e.g. Pesonen et al., 2003). Furthermore, global peaks in isotopic age determinations at 2.7-2.6 Ga,
1.9-1.8 Ga and 1.2-1.1 Ga coincide with the existence of supercontinents prior to Pangea and
Gondwana (Reddy and Evans, 2009) namely Rodinia (ca. 1.05-0.7 Ga), Nuna (ca. 1.8-1.26 Ga) and
Kenorland (2.7-2.1 Ga) (Pesonen et al., 2012b).
Gurnis (1988) proposed a supercontinent model, in which supercontinents insulate the mantle beneath
them, eventually leading to a breakup over a geoid high, which is created by rising plumes. The rising
plumes/superplumes in turn are proposed to exist as a result of slab avalanches caused by slabs of
subducted oceanic lithosphere that drop into the lower mantle (e.g. Condie, 1998; 2001). The
subsequent development of supercontinents after break-up can be explained by three different models,
namely extroversion, introversion and orthoversion (e.g. Murphy and Nance, 2013; Mitchell et al.,
12
2012). Extroversion refers to a process in which dispersing continents move towards subduction
zones in the exterior ocean where they amalgamate and form a new supercontinent. Introversion, as
was the case during the amalgamation of Pangea, involves a reversal in the direction of the motion of
continents leading to the subduction of the interior oceanic crust resulting in a new supercontinent
where its predecessor was located. Orthoversion refers to the process in which a new supercontinent
forms at a location orthogonal to from its predecessor and involves the subduction of either exterior
or interior oceanic crust, or a combination (Murphy and Nance, 2003; Mitchell et al., 2012).
Supercontinents have played a key role in the evolution of the Earth’s surface at least since
Paleoproterozoic time (Evans and Pisarevsky, 2008). For example the compression of continents
associated with supercontinent amalgamation is expected to lower sea levels, and crustal thinning and
seafloor spreading associated with supercontinent break-up is expected to give rise to an increase in
sea level (Reddy and Evans, 2009 and references therein). The increase or decrease in energyabsorbing oceanic surfaces may affect global climate. Supercontinent cycles also have significant
effects on the carbon cycle, as higher precipitation rates on low-latitude supercontinents enhance
silicate weathering (e.g. Kirschvink, 1992).
1.1.1. Nuna
Orogenic activity between 2.1 Ga and 1.6 Ga (Zhao et al., 2004, Rogers and Santosh, 2004) and
evidence of widespread extension starting at ca. 1.5 Ga (Rogers and Santosh, 2004) indicates the
existence of the pre-Rodinian superconinent Nuna (also known as Columbia or Hudsonland). Several
reconstructions of Nuna have been proposed (e.g. Rogers, 1996; Hoffman, 1997; Meert, 2002; Zhao
et al., 2004, Pisarevsky et al., 2014a), most of which feature Laurentia, Baltica, Amazonia, Siberia,
and Australia connections.
The core of the supercontinent Nuna, NENA (North Europe – North America), in which Greenland
aligns with Kola Peninsula and northern Norway, is proposed by many authors to have existed during
1.8-1.2 Ga (e.g. Gower et al., 1990; Buchan et al., 2000; Pesonen et al., 2003; Salminen and Pesonen,
2007; Evans and Pisarevsky, 2008; Evans and Mitchell, 2011; Salminen et al., 2014). It is supported
by a tight cratonic fit, concordance with basement geology, geochronological correlations, and
matching paleomagnetic poles at 1.75 Ga, 1.63 Ga, 1.59 Ga, 1.46 Ga and 1.27 Ga (Salminen et al.,
2014).
13
Baltica consists of the Sarmatian craton, Volgo-Uralian crustal segment, and Fennoscandian Shield.
Sarmatia and Volgo-Uralia collided at ca. 2.1-2.0 Ga forming Sarmatia/Volgo-Uralia, which later
collided with the Fennoscandian Shield between 1.82 and 1.80 Ga along the Russian collisional belt
(Bogdanova et al., 2008, 2015; Elming et al., 2010).
Based on bedrock geological studies, Johansson (2009) proposed a PaleoproterozoicMesoproterozoic assembly between Baltica and Amazonia (SAMBA model) as the 1.9-1.8 Ga
Svecofennian orogen in Baltica continues into the 1.98-1.81 Ga Ventuari-Tapajós province in
Amazonia. The slightly younger 1.85-1.65 Ga Transscandinavian igneous belt (TIB) and 1.64-1.52
Ga Gothian orogens of Baltica continues into the 1.78-1.55 Ga Rio Negro-Juruena province in
Amazonia. It is widely accepted that Amazonia and West Africa formed a rigid continent, similar to
their Gondwana configuration, during the Mesoproterozoic (Trompette, 1994). Some authors (e.g.
Pisarevsky et al., 2014a) position Amazonia together with West Africa, in the antiparallel posision,
as an alternative to the SAMBA configuration.
Laurentia consists of Superior, Wyoming, Slave, Nain, and Churchill (Hearne, Rae and Burwell
domains) cratons which were assembled by collisional orogens by 1.8 Ga (Buchan et al., 2000;
Hoffman, 1989). Slave and Rae collided at 1.97 Ga (Thelon orogeny), and the the Slave-Rae-Hearne
assembly collided with Superior between 1.91 and 1.81 Ga, forming the Trans-Hudson orogeny
(Hoffman, 1988; Mitchell et al., 2014).
It is proposed that Siberia also formed part of the core of Nuna (Wingate et al., 2009; Evans and
Mitchell, 2011), since the Siberian APWP can be juxtaposed with that of NENA between 1.8 Ga and
1.4 Ga (Evans and Mitchell, 2011). Geological correlations suggest that the Aldan shield of Siberia
was connected to northern Greenland and Canada (Rainbird et al., 1998). Pisarevsky et al. (2014a)
on the other hand proposes a distant, but fixed position of Siberia relative to Laurentia between 1.5
Ga and 1.0 Ga, based on coeval paleomagnetic poles (Wingate et al., 2009; Meert and Stuckey, 2002).
Australia formed by the assembly of the North Australian, West Australian and South Australian
cratons. The time of assembly is uncertain, however Li and Evans (2011) proposed that they were in
close proximity to each other since ca. 1.8 Ga. North Australia was possibly connected to NW
Laurentia during the Mesoproterozoic (Zhang et al., 2012; Pisarevsky et al., 2014a).
Mafic magmatism of ca. 1.59 – 1.46 Ga across Nuna, and dyke swarms of the 1.38-1.24 Ga period,
are linked with the break-up of Nuna (Roberts, 2013 and references within). However, many of the
dyke swarms occurring at craton margins may be related to plate-margin processes and not the breakup of the supercontinent (Söderlund et al., 2005; Shellnutt and MacRae, 2012; Roberts, 2013).
14
Furthermore, the break-up of supercontinents are followed by rifting and the gradual formation of
passive margins, however Nuna has a low passive margin abundance, suggesting that no significant
break-up and dispersal of continents took place (Bradley, 2008; 2011).
Fig. 1. Examples of Nuna and Rodinia reconstructions.
1.1.2. Rodinia
Rodinia assembled through orogenic activity between 1.3 Ga and 0.9 Ga (e.g. Li et al., 2008), and
reached complete assembly at ca. 1.0 Ga (Rogers and Santosh, 2004). All continental blocks were
likely to be involved and Laurentia is thought to have formed the core of Rodina as it is surrounded
by Neoproterozoic passive margins (McMenamin and McMenamin, 1990; Hoffman, 1991; Meert and
Torsvik, 2003; Li et al., 2008). Different configurations of Rodinia have been proposed, including
the SWEAT, AUSWUS, AUSMEX, and “missing-link” models, each containing difficulties. The
SWEAT (Southwest U.S. – East Antarctica) model (Fig. 1), proposed by Moores (1991), and shown
in Fig. 1, assumes that Grenville orogenic belts are zones of oceanic closure between continental
blocks (Rogers and Santosh, 2004). In this configuration eastern Antarctica and eastern Australia are
connected to southwestern Laurentia. The AUSWUS (Australia-Southwest U.S.) model is based on
15
matching basement provinces and sedimentary provenances between Australia and southwestern
Laurentia (Karlstrom et al., 2001; Burrett and Berry, 2000). Based on paleomagnetic results, Wingate
et al. (2002) proposed the AUSMEX (Australia – Mexico) model in which northeastern Australia is
joined to Mexico of southern Laurentia. The “missing link” model (Li et al., 1995) proposes that the
South China block was situated between Australia-East Antarctica and Laurentia due to the
mismatches in crustal provinces of Australia-East Antarctica and Laurentia in the SWEAT model as
well as similarities in stratigraphy and crustal provinces between South China, southeastern Australia
and western Laurentia.
Meert (2014a) points out the similarities between Nuna, Rodinia and Pangea, in particular the
assembly of Baltica, Laurentia and Siberia, as well as parts of East Gondwana. Such recurring
similarities would be unlikely with extroversion and introversion of ocean basins leading Meert
(2014a) to suggest lid-style tectonics with episodes of true polar wander (i.e. the rotation of the Earth
with respect to its spin axis). Roberts (2013) proposes a model in which Nuna underwent only minor
modifications to form Rodinia at 1.1-0.9 Ga, which include the transformation of external
accretionary belts into the internal Grenville and equivalent collisional belts.
Piper (1982) proposed an alternative Proterozoic supercontinent, namely Palaeopangea, that are also
based on lid-style tectonics, and survived until a major break-up event during the Ediacaran (ca. 0.6
Ga; Piper, 2013). Piper (1982, 2013, 2015) further showed that the paleomagnetic record conformed
to a single, global APWP using the Paleopangea reconstruction from ca. 2.7 Ga to 0.6 Ga and that
the majority of paleomagnetic poles at 2.7-2.5 Ga, 1.5-1.2 Ga, and 0.75-0.6 Ga are positioned in a
quasi-static position. Palaeopangea is a crescentric shaped supercontinent, similar in shape as Pangea,
confined to a single hemisphere (Piper, 2015). Palaeopangea is made up of the 3.0 Ga Ur (Kaapvaal,
Pilbara, Dharwar and Signhbum cratons), 2.5 Ga Arctica (Siberian and Canadian cratons), and 2.0
Ga Atlantica (South American and West African cratons). The period 2.1-1.8 Ga was marked by
significant crustal growth forming long linear mobile belts within the core of Ur-Atlantica-Arctica
and accreted to the periphery, with the orogenic belts forming an axial trend through the continental
lid (Piper, 2015). Similarly, the Grenville orogenic belts (1.1 Ga) formed linearly along Palaeopangea.
However, the Palaeopangea model is critized for being based on incorrect application of
paleomagnetic data, and for lacking geological evidence (Li et al., 2009).
16
1.2. Paleomagnetism
1.2.1. The geocentric axial dipole
The role of paleomagnetism in supercontinent reconstruction is based on the concept of a geocentric
axial dipole (GAD) model. According to this model the Earth’s magnetic field is produced by a single
magnetic dipole (consisting of a North and South pole) at the centre of the Earth, and is aligned with
the Earth’s rotation axis (Butler, 2004). The magnetic field vector (H), at any one location on Earth,
is expressed:
‫ܪ‬ൌ
‫ܯ‬
ඥͳ ൅ ͵‫݊݅ݏ‬ଶ ߣ
ܴ௘ଷ
(1)
And can be divided into a vertical (Hv) and horizontal (Hh) component:
‫ܪ‬௩ ൌ
ʹ‫ߣ݊݅ݏܯ‬
ܴ௘ଷ
(2)
‫ߣݏ݋ܿܯ‬
ܴ௘ଷ
(3)
‫ܪ‬ℎ ൌ
In Eq. (1) – (3) M is the dipole moment of the geocentric axial dipole (Am2), λ is the geographic
latitude (ranging from -90° to +90°), and Re is the mean radius of the Earth (6371 km). The magnetic
field strength observed at the Earth’s surface increases from the equator to the poles. The magnetic
inclination (I) is defined as the vertical angle between the horizontal component (Hh) and the total
magnetic field vector (H) for a given latitude. The relationship between magnetic inclination and
geographic latitude (λ) is expressed by the following equation:
tanI = 2 tanλ
(4)
where I is magnetic inclination and λ is the geographic latitude.
17
Fig. 2. Geocentric axial dipole model (from Butler (2004)), illustrating the relationship between latitude (λ) and magnetic
inclination (I).
The reconstruction of Precambrian supercontinents using paleomagnetism assumes that 1) the Earth
has had an internal magnetic field since at least the Paleoproterozoic, and 2) that the Precambrian
field was generated by a GAD. Paleosecular variation studies support a geodynamo during
Neoarchean and Proterozoic times (e.g. Smirnov et al., 2011). Inclination frequency analysis
(Veikkolainen et al. 2014a), geomagnetic field reversal symmetry studies (Veikkolainen et al.
2014b), and paleointensity investigations (Biggin et al., 2015), indicate a predominantly GAD field,
with weak axial quadrupole and octopole coefficients (<10%) that have negligible consequences to
Precambrian continental assemblies. Furthermore, Precambrian magnetostratigraphy (Pavlov and
Gallet, 2011) and latitudinal control of paleoclimatic indicators, such as evaporates (mid-latitudes) or
glacial deposits (polar latitudes) further support a GAD (Evans, 2006).
1.2.2. Magnetism in rocks
Recording the magnetic field vector of a rock formation depends on the material. Induced
magnetization is linearly related to the magnetic field and is expressed:
Ji = χH
(5)
where χ is the magnetic susceptibility of the material (dimensionless, or ‘SI’), and H is the magnetic
field (Am-1). Rocks contain diamagnetic, paramagnetic, and ferromagnetic minerals. Ferromagnetism
is used here in a broad sense, and unless indicated otherwise, also implies ferri- and
antiferromagnetism.
18
In response to a magnetic field, diamagnetic particles, such as quartz, acquire a small induced
magnetization opposite to the applied field. Diamagnetic magnetization is very small, does not depend
on temperature, and becomes zero in the absence of an external magnetic field. Paramagnetic
minerals, such as biotite and pyrite, acquire an induced magnetization parallel to the magnetic field.
Paramagnetic magnetization is relatively small and depends on temperature. In the absence of an
external magnetic field the magnetization is zero (Dunlop and Özdemir, 1997).
Only ferromagnetic particles have the ability to record the direction of the applied magnetic field so
that it exists in the absence of the applied magnetic field. This is known as remanent magnetization,
which is the focus in paleomagnetic research. Unlike dia- and paramagnetic magnetization,
ferromagnetic materials exhibit hysterisis (an irreversable path of magnetization) in response to an
external magnetic field. Ferromagnetic elements also acquire induced magnetizations, which are
orders of magnitude higher than paramagnetic minerals, due to strong interactions between spin
moments.
Natural remanent magnetization (NRM) of a rock is obtained from the geomagnetic field during rock
formation and the subsequent geological history of the rock. NRM can be stable over geological time
and resist the changes of the Earth’s magnetic field. Although remanent magnetization undergoes a
natural viscous decay with time, described as its relaxation time, at typical Earth surface temperatures
the relaxation time for ferromagnetic grain sizes ≤100μm is of the order of 109 yr or more (Butler,
2004).
There are different types of natural remanent magnetization, depending on how it is acquired.
Thermoremanent magnetization (TRM) is acquired during cooling after formation (igneous rocks) or
after high metamorphic temperatures. If the metamorphic temperature is below the original formation
temperature, the rock may acquire a partial thermoremanent magnetization (pTRM). Each given
ferromagnetic mineral has a characteristic Curie temperature, above which the mineral becomes
paramagnetic. Relaxation time is temperature dependent. As the rock cools below its Curie
temperature, the relaxation time increases rapidly with decreasing temperature. For a given relaxation
time there is a specific temperature known as the blocking temperature. As the ferromagnetic grains
cool below their respective Curie temperatures and the temperature decreases through the blocking
temperatures the grains acquire a TRM along the prevailing field, and experience a significant
increase in relaxation time, becoming increasingly stable. Table 1 shows the properties of different
ferromagnetic minerals. Magnetite, a ferrimagnetic iron oxide, does not only have a relatively high
Curie temperature (578°C) but also strong magnetic susceptibility (χ) as well as high saturation
19
magnetization (Js) and high to moderate coercivity values, making it one of the most important
minerals in igneous rocks for paleomagnetic studies.
Table 1
Chemical composition and basic magnetic properties of the main ferromagnetic minerals (Dunlop
and Özdemir, 1997).
Mineral
Composition
TC (°C)
χ (SI)
JS (kA/m)
HC (T)
Goethite
αFeOOH
120
ca. 20-30*
ca. 2
2.2
Magnetite
Fe3O4
580
1.2-19.2
480
0.1-0.25
Maghemite
γFe2O3
590-675
**
380
ca. 0.1
Hematite
αFe2O3
675
0.5-35 × 10-3 ca. 2.5
2.5-7.6
Titanomagnetite
Fe3-xTixO4
150-540
***
125
0.2
Pyrrhotite
Fe7S8
320
0.001-6
ca. 80
0.2-0.35
TC = Curie temperature, χ = magnetic susceptibility, JS = saturation magnetization, HC = coercivity
* Özdemir and Dunlop (1996)
** Difficult to measure in laboratory due to chemical alteration
*** Depends on Ti-content: decreases with increasing Ti-content.
Chemical changes of a rock can give rise to chemical remanent magnetization (CRM) or thermochemical remanent magnetization (TCRM) if the chemical changes involve a rise in temperature.
CRM occurs as a result of the alteration of pre-existing minerals to a ferromagnetic mineral or the
precipitation of ferromagnetic minerals from a solution (Butler, 2004). CRM is common in
sedimentary rocks. In marine and lake sediments CRM can form soon after deposition as as ironsulfate reduction diagenesis (Kodama, 2012). When secondary magnetic minerals first form they are
too small to carry a stable remanent magnetization, however as they grow they can eventually attain
stable remanance with relaxation times of the order of 109 yr.
Detrital remanent magnetization (DRM) is acquired during deposition and lithification of
sedimentary rocks. During deposition pore spaces in the non-magnetic matrix are larger than the
magnetic grains. The magnetic grains can therefore freely move and align with the Earth’s magnetic
20
field. As the sediment is buried and becomes dewatered from compactions, the pore spaces become
smaller and the magnetic minerals become trapped and immobile obtaining a post-depostional
remance (pDRM). The pDRM is usually aquired at a burial depth of below 20 cm (eg. Tauxe et al.,
1996; Kodama, 2012).
Viscous remanent magnetization (VRM) is gradually acquired when a formation is exposed to weak
magnetic fields, like the Earth’s magnetic field, and typically has low unblocking temperatures. The
direction of VRM gradually aligns with the ambient magnetic field direction. Characteristic remanent
magnetization (ChRM), on the other hand, is the highest stability component of a sample during the
course of demagnetization, and is usually an ancient magnetization. ChRM can be TRM, CRM,
TCRM, or DRM (Butler, 2004).
1.2.3. Paleomagnetic measurements
Remanent magnetization of a rock can be determined through paleomagnetic measurements. By
means of demagnetization methods, secondary NRM is removed and the ChRM is isolated. The two
most common demagnetization techniques are alternating field (AF) demagnetization and thermal
demagnetization.
During alternating field demagnetization the specimen is exposed to an alternating magnetic field in
a zero ambient field. The applied alternating field decreases in magnitude with time, magnetizing all
ferromagnetic minerals with a coercivity smaller than the peak demagnetizing field in alternating
antiparallel directions, resulting in a net magnetization of zero. Coercivity is defined as the intensity
of the applied magnetic field required to reduce the magnetization of the material to zero. The
stepwise alternating field process is repeated with a progressively increasing peak applied field until
the specimen is demagnetized completely, or the maximum field of the instrument is reached. A
typical AF progression would comprise peak fields of 2.5, 5, 7.5, 10, 15, 20, 25, 30, 40, 50, 60, 80,
100, 120, 140, and 160 mT.
Thermal demagnetization involves the heating and subsequent cooling to room temperature of a
specimen in zero magnetic field. This is repeated with progressively increasing temperatures, and
with each step the remanent magnetization with a blocking temperature equal to or lower than the
applied temperature is demagnetized. A typical thermal progression would be temperatures of 200,
300, 400, 500, 520, 540, 560, 570, 580, 600, 620, 640, 660, 670, and 680°C, tailored to reveal the
21
characteristic unblocking temperature ranges of magnetite (500° - 580°C) and hematite (620° 680°C)
After each AF or thermal step, the remaining remanent magnetization of the vertical and horizontal
axes (Mx, My, Mz) is measured, and from these the total magnetic moment (M), and magnetic
declination (D) and inclination (I) are calculated as follows:
‫ ܯ‬ൌ ට‫ܯ‬௫ଶ ൅ ‫ܯ‬௬ଶ ൅ ‫ܯ‬௭ଶ
‫ ܦ‬ൌ ‫ି݊ܽݐ‬ଵ ൬
‫ ܫ‬ൌ ‫ି݊ܽݐ‬ଵ ቆ
(6)
‫ܯ‬௬
൰
‫ܯ‬௫
‫ܯ‬௭
ඥ‫ܯ‬௫ଶ ൅ ‫ܯ‬௬ଶ
(7)
ቇ
(8)
To isolate different remanent components, linear least square analysis (Dunlop, 1979) is a commonly
used method. For visualization of remanence components, a Zijderveld diagram (Zijderveld, 1967) is
used to identify as well as quantify multiple components. A Zijderveld diagram is a plot of
magnetization direction and intensity during the course of AF or thermal demagnetization, projected
on two orthogonal planes. The slopes of linear successions of demagnetization steps, in three
dimensions, corresponds with the declination and inclination of magnetic components. Other methods
for quantifying remanence components include principal component analysis (Kirschvink, 1980),
intersecting great circle technique, and analysis of vector end points (Halls, 1976; 1978). Fisher
statistics are used to calculate mean declination and inclination values for components of samples or
sites (Fisher, 1953)
1.2.4. Apparent polar wander paths and paleogeographic reconstructions
Paleomagnetic pole positions are calculated from paleomagnetic declination and inclination. A
paleomagnetic pole is the position of the North pole at the time of magnetization in relation to the
current geographic position of the craton/continent under investigation. The position of the
paleomagnetic pole is independent of the observing locality, and is therfore used to compare
paleomagnetic results between different localities or continents. Paleomagnetic poles are used to
construct apparent polar wander paths (APWP's). An APWP is a series of age-successive
paleomagnetic poles indicating the movement of a continental/craton over time. APWP’s from
22
separate continents/cratons can be used to test their assembly and break-up. The denser the data set
of reliable paleomagnetic poles for each APWP, the more precisely the periods of amalgamation and
break-up between continents/cratons can be determined (Evans and Pisarevsky, 2008).
In articles I-IV, the most reliable poles were used for any given time to construct APWP’s. The
reliability of paleomagnetic poles is determined by seven criteria proposed by Van der Voo (1990).
They are i) a well determined rock age that corresponds with the age of magnetization, ii) a sufficient
number of samples (N>24), iii) adequate demagnetization (including vector subtraction), iv) a
positive field stability test (such as a baked contact test), v) tectonic coherence with the
craton/continent in question, vi) the presence of anti-parallel polarities, and vii) no resemblance to
younger paleopoles. The more of these criteria a paleomagnetic pole fulfills the more reliable it is,
however for it to be regarded a key paleomagnetic pole (i.e. well dated and well defined pole for
reconstruction) it has to fulfill at least the basic criteria: a well determined rock age and a field stability
tests, in addition to other criteria (Buchan et al. 2000; Buchan, 2013).
Paleomagnetism is an essential tool in paleogeographic reconstructions to show the ancient positions
of continents/cratons relative to each other. Combined with geochronology, paleomagnetism remains
the only quantitative method to reconstruct Precambrian continents to an absolute paleogeographic
reference frame. The paleolatitude of a continent is determined by the magnetic inclination. Magnetic
declination determines the vertical axis rotation of the continent. Paleolongitude cannot be determined
with paleomagnetism, so geological features such as dyke swarms and orogenic belts are used to
constrain the relative positions of continental blocks.
1.2.5. Mafic dykes and sedimentary rocks in paleomagnetic studies
Mafic dykes are ideal for paleomagnetic studies, because they cool rapidly and therefore provide an
accurate record of the Earth’s magnetic field. In addition, they may contain zircon and/or baddeleyite
which are used to establish the precise crystallization ages of the dykes (e.g. French and Heaman,
2010; Halls et al., 2007). Sedimentary rocks, on the other hand, present challenges as they are more
difficult to be dated accurately, and they can be subject to inclination shallowing. However,
sedimentary rocks can give a continuous record of the magnetic field (Kodama, 2012) and therefore
fill important age gaps in paleomagnetic records. Another reason why the continuity of sedimentary
paleomagnetic records is important is because it allows time averaging of the secular variation of the
geomagnetic field (Butler, 2004; Kodama, 2012).
23
It is important to make inclination corrections for paleomagnetic results obtained from sedimentary
rocks, since inclination shallowing can be as much as 20° at mid-latitudes (Kodama, 2012).
Inclination shallowing occurs as a result of burial compaction at depths of tens to hundreds of meters
(Kodama, 2012). Examples of sedimentary inclination shallowing are numerous in a variety of natural
settings (e.g. Bilardello and Kodama, 2010) and in several laboratory re-deposition experiments (e.g.
Tauxe and Kent, 1984). Due to the great diversity of sediment characteristics and depositional
conditions, in addition to the assortment of specific mechanisms by which a magnetization may
acquire a shallow inclination bias, the magnitude of the bias in sedimentary rocks varies (Domeier et
al., 2012). There are several ways to correct this including: 1) the DRM tensor approach of Jackson
et al. (1991); 2) the high field anisotropy of isothermal remanent magnetization (hf-AIR) approach
of Bilardello and Kodama (2009); 3) the Elongation/Inclination (E/I) method of Tauxe and Kent
(1984) using the statistical field model TK03.GAD; 4) the E/I method of Tauxe and Kent (2004)
using samples from a single horizon and assuming a Fisher (1953) distribution; and 5) a blanket
correction based on assumed values for flattening factor f using an inclination correction equation by
King (1955):
tan(Io) = f tan(If)
(9)
where Io is the observed inclination, If is the field inclination and f is an empirically derived ‘flattening
factor’ (Tauxe et al., 2008). According to the compilation of Bilardello and Kodama (2010), f-values
from magnetite dominated sedimentary rocks range from 0.54 to 0.79, with a mean of 0.65, whereas
hematite-dominated sedimentary rocks have yielded f values from 0.4 to 0.83, with a mean of 0.59.
The basic requirement in making paleomagnetic reconstructions is to verify that the observed
magnetization direction is original, recording a primary ambient field direction during emplacement
of the rock formation. In this study a baked contact test, and a fold test were applied. In the case of
remanent magnetization of thermal origin it can be tested by the baked contact test (Everitt and Clegg,
1962). The concept of the test is that an igneous body, e.g. a dyke, heats (or bakes) the contact zone
of the host rock to near or above the Curie temperature of the magnetic minerals. Such minerals gain
the direction of the Earth’s magnetic field at the cooling time. This direction is likely to be similar to
the remanence direction in the dyke itself as both cool simultaneously. As temperature decreases with
increasing distance from the dyke, the remanence directions of the baked zone gradually change. A
positive baked contact test therefore requires that the dyke and its baked host rock carry similar stable
remanence directions, and that the unbaked host rock carries a significantly different stable direction
24
(e.g. Irving, 1964; Buchan, 2013). The baked contact test is strengthened if the host rock contains a
hybrid zone, where both host rock and dyke remanent components occur together in addition to the
two above mentioned remanence zones (Buchan, 2013).
In the case of sedimentary rocks, the fold test (e.g. McFadden, 1990), establishes that the age of
remanence is pre-, syn-, or post-fold. In the fold test, if a ChRM was acquired prior to folding,
directions of ChRM from different limbs of a fold are dispersed, but they converge when the structural
correction is made. If the structural correction results in more dispersed ChRM direction, the ChRM
direction was acquired after folding took place. Unless it can be demonstrated that folding occurred
during deposition of the rock unit under study, it does not establish the primary nature of remanent
magnetization. This can be established if the maximum clustering of ChRM directions occurs at less
than 100% unfolding.
2. Aims of this study
The main objective of this work was to obtain new Precambrian paleomagnetic data for Baltica, to
refine the Precambrian APWP of Baltica, and to test the assemblies of the Precambrian
supercontinents Nuna and Rodinia.
The objective with studying the Keuruu diabase dykes was to establish the position of Baltica related
to Laurentia at the onset (1.87 Ga) of the NENA assembly. With the Åland intrusives and Satakunta
sandstone formation the objective was to test the proposed 1.8-1.2 Ga NENA connection within the
supercontinent Nuna. In the case of the Hailuoto sedimentary formation, the objective was to test the
Late Neoproterozoic paleogeographic positions of Baltica to better understand the environmental
conditions during that time. With a new Late Neoproterozoic pole a further aim was to explore if the
large sways in the Ediacaran APWP can be explained by high plate velocities alone.
In order to achieve firm conclusions an additional aim of this study was to obtain precise U-Pb ages
from Åland and Keuruu diabase dykes, and to obtain a new relative paleomagnetic age of deposition
from the paleomagnetic results of Satakunta sandstone.
3. Materials and methods
3.1. Sampling and sample preparation
Field sampling for this work was carried out on the Satakunta sandstone formation, Åland diabase
and quartz porphyry dykes, and the Keuruu diabase dyke swarm. The samples were taken with a
portable rock drill or as oriented block samples. The samples were oriented with a magnetic and/or
25
sun compass and the local declination was corrected. Multiple specimens were prepared from each
sample; block samples were drilled in the laboratory and all the drill cores were sawn into standard
cylindrical specimens (2.56 cm diameter (D), ~2.3 cm length (h), with h/D ≈ 0.9). Further samples
were obtained from an inclined 272 m deep, fully oriented sedimentary drill core from the Hailuoto
sedimentary formation. Samples with a maximum length of about 6 cm were first cut from the drill
core and from these 7 - 8cm3 cubic specimens parallel to the core axis, were prepared.
3.2. Laboratory measurements.
Measurements were carried out at the Solid Earth Geophysics Laboratory of the University of
Helsinki (UH), Finland (Pesonen, 2006) and at the Paleomagnetic laboratory of the Geological Survey
of Finland (GTK), Espoo (Pesonen, 1989).
3.2.1. Petrophysical measurements
Magnetic susceptibility, intensity of natural remanent magnetization (NRM), and rock density were
measured for each specimen using the Risto-5 petrophysics assembly of the Solid Earth Geophysics
laboratory at the University of Helsinki. The magnetic susceptibility was measured at room
temperature with the Risto-5 kappabridge at a frequency of 1025 Hz and a field of 48 A/m and the
NRM was measured with the Risto-5 fluxgate magnetometer. From the NRM and magnetic
susceptibility the Königsberger (Q) ratio was calculated. The Q ratio is the ratio of remanent
magnetization to the induced magnetization in Earth’s magnetic field using the formula:
Q = μ0NRM/χB
(10)
Where μ0 is the magnetic permeability in a vacuum (4πx10-7H/m), NRM is the natural remanent
magnetization (A/m), χ is magnetic susceptibility (SI), and B is the geomagnetic reference field value
(for Finland a value of 50 μT is used). If Q>1, the remanent magnetization is stronger than the induced
magnetization and the specimen is able to maintain a stable remanence. Specimens with high Q values
(>10) must be treated with caution, since it may indicate either fields other than the Earth’s internal
magnetic field (e.g. lightning induced fields), or unstable magnetic carriers.
26
3.2.2. Paleomagnetic and rock magnetic measurements
Stepwise alternating field (AF) demagnetizations were done using a Schonstedt SD-1 demagnetizer
up to 100 mT (GTK) or a three-axis demagnetizer with maximum field up to 160 mT, coupled with
a 2G–DC SQUID magnetometer (UH) or a 2G-RF SQUID magnetometer (GTK). Stepwise thermal
demagnetizations were performed using a GTK built or Schonstedt TSD-1 furnace (GTK), an
MMTD1 thermal demagnetizer (UH), or an ASC Model TD48-SC controlled atmosphere (Ar)
furnace (UH). Remanent magnetizations were measured with a 2G–DC (or 2G–RF) cryogenic
magnetometer or a GTK-built spinner magnetometer (Pesonen et al., 1989).
Magnetic carriers were studied by thermomagnetic analysis of selected specimens using an Agico
CS3-KLY-3S Kappabridge system (UH), which measures the bulk susceptibility (k) of the samples
during heating from -192° to 700°C and cooling back to room temperature (in Argon gas).
Ferromagnetic minerals (senso lato) are identified by changes in magnetic susceptibility with
increasing temperature.
Vector components were visually identified using stereographic and orthogonal projections
(Zijderveld, 1967) and the directions were calculated by a least squares method (Leino, 1991). Mean
remanence directions for the different components were calculated according to Fisher (1953), giving
a unit weight to each sample (each specimen has a unit weight within a sample) to compute site mean
directions and corresponding paleomagnetic poles (Irving, 1964). The paleogeographical
reconstructions were done with GMAP (Torsvik and Smethurst 1999), and GPlates programs.
Inclination corrections were made for paleomagnetic results obtained from sedimentary rocks. In
Article III we used Eq. 9 to correct inclination shallowing of the Hailuoto sedimentary rocks. In
Article I, high field anisotropy of isothermal remanent magnetization (hf-AIR) measurements were
carried out to correct inclination shallowing of the Satakunta sandstone. The hf-AIR measurement
technique of Bilardello and Kodama (2009) was followed; 5T isothermal remanent magnetization
(IRM) was applied in 9 orientations (Girdler, 1961) of oriented samples, using a MMPM 10 pulse
magnetizer. Samples were then thermally demagnetized at 125°C in a MMTD1 thermal
demagnetizer, to demagnetize goethite. The samples were then AF demagnetized at 100 mT, to
demagnetize magnetite. The remaining hematite component was measured in a 2G SQUID
magnetometer. The best-fit anisotropy tensor was determined using the least square fitting, and the
eigenvalues were calculated from the tensor. The inclination correction was applied by using the
hematite correction equation of Tan and Kodama (2003):
27
Tan Ic = tan Im (2a+1)kmax-1 / (2a+1)kmin-1
(11)
where Ic and Im are the corrected and measured inclinations, respectively, kmax and kmin are the
normalized maximum and minimum principal axes, respectively of the fabric, and a is the individual
particle anisotropy. The latter was not measured, but a value of 1.4 was used, which is considered to
be reasonable for hematite (Tan et al. 2007; Kodama, 2012).
In addition to the inclination correction, a tilt correction was carried out on the Satakunta sandstone
samples (Article I). Baked contact tests were carried out on Åland (Article II) and Keuruu (Article
IV) dykes of both polarities.
4. Review of articles
4.1. Article I
Klein, R., Pesonen, L.J., Salminen, J., Mertanen, S., 2014. Paleomagnetism of Mesoproterozoic
Satakunta sandstone, Western Finland. Precambrian Research 244, 156-169.
In this article a paleomagnetic study of the Mesoproterozoic Satakunta sandstone formation is
presented. The aim of the study is to identify the ChRM of the Satakunta sandstone in order to
estimate the relative paleomagnetic age of deposition, and to use the new paleomagnetic pole for
Baltica to check the validity of NENA at the estimated age of the sandstone.
The Satakunta sandstone formation is a fluvial formation situated in the Satakunta region in Western
Finland (61.25°N; 22°E). The formation consists of unmetamorphosed sandstone that was deposited
in a deltaic environment and has been preserved in a SE-NW trending graben within the
Paleoproterozoic Svecofennian Domain. It is part of a much wider sandstone basin extending across
the Baltic Sea to the Gävle valley and Nordingrå area in Sweden (Kohonen et al., 1993; Pokki et al.
2013).
Two components of natural remanent magnetization (NRM) were isolated in the sandstone with
alternating field (AF) and thermal demagnetization treatments. The first is a high
coercivity/unblocking temperature component which, after tilt and inclination shallowing corrections,
yields a dual polarity remanent magnetization of D = 25.2°, I = 3.9° with α95 = 9.1° (10 sites),
corresponding to a paleomagnetic pole of Plat. = 27.8° N, Plon = 173.2° E, with A95 = 6.5°. Positive
reversal and tilt tests, an inverse baked contact test, as well as a stratigraphic order of the site mean
poles, indicates a primary remanence. The second component is a lower coercivity/unblocking
28
temperature component with a remanent magnetization direction D = 34.0°, I = - 40.7° with α95 =
8.1° (13 sites) similar to the 1.26 Ga Postjotnian Satakunta diabase intrusions, which are widespread
in the area. This indicates that the component reported as the Satakunta sandstone component in an
earlier study by Neuvonen (1973), is an overprint (baked) component.
The relative age of the sandstone is estimated at ca. 1600 based on the proximity of the Satakunta
sandstone pole to other well defined poles of 1452-1630 Ma from Baltica. Comparing APWPs of
Baltica and Laurentia during 1.77-1.27 Ga gives strong support to the NENA configuration during
the Mesoproterozoic.
4.2. Article II
Salminen, J.M., Klein, R., Mertanen, S., Pesonen, L.J., Frödjö, S., Mänttäri, I., Eklund, O. 2015.
Palaeomagnetism and U–Pb geochronology of c. 1570 Ma intrusives from Åland archipelago, SW
Finland –implications for Nuna. In Li, Z. X., Evans, D. A. D. and Murphy, J. B. (eds) Supercontinent
Cycles Through Earth History. Geological Society, London, Special Publications 424, doi:
10.1144/SP424.3
In this article, new paleomagnetic and isotope age data of Early Mesoproterozoic (i.e. Subjotnian)
intrusives from the Åland archipelago, SW Finland is reported. The motivation for this study was to
obtain Early Mesoproterozoic key-paleomagnetic data for Baltica, to obtain a preliminary pole for
the rapakivi granite in Åland, to refine the Early Mesoproterozoic APWP of Baltica, to obtain an
additional precise U–Pb age from a reversely magnetized dyke in Åland (Korsö) and finally to test
the proposed NENA connection within supercontinent Nuna.
The Åland archipelago is situated at the entrance to the Gulf of Bothnia in the Baltic Sea. In Åland,
the Svecofennian crust consists of Paleoproterozoic granitoids, paragneisses, and felsic and
intermediate metavolcanic rocks. These are cut by Subjotnian rapakivi granite intrusions, mafic
dykes, as well as quartz porphyritic dykes generally trending SSW–NNE. Several Subjotnian units in
Åland have been dated before. The rapakivi granites yield U–Pb (zircon) ages of 1568±10 Ma and
1579±13 Ma (Suominen 1991), correlating with ages from Vehmaa rapakivi granite in mainland
Finland, NE from Åland. Previous U–Pb (zircon) ages for diabase dykes in Åland vary between
1540±12 Ma and 1577±12 Ma. Previous U–Pb (zircon) ages for quartz porphyry dykes range between
1576±13 Ma, 1576±13 Ma and 1571±20 Ma (Suominen 1991).
29
The paleomagnetic results reveal dual-polarity magnetizations with a pronounced reversal
asymmetry. The asymmetry is explained by an unremoved secondary component, which is affecting
the N-polarity dykes more. The primary nature of magnetization in dykes for both normal (N) and
reversed (R) groups is verified by positive baked contact tests. A dyke showing reversed polarity
from Korsö is dated 1575.9±3.0 Ma (U–Pb, zircon) in this study. This and previous U–Pb data tighten
the magmatic activity in Åland to 1580–1570 Ma. The new paleomagnetic data is combined with
those from earlier studies to provide a new key paleomagnetic pole for Baltica. The normal polarity
ChRM (D = 12.4°; I = 10.5°; α95 = 6.9°; k = 12.9; N = 31) yields a paleomagnetic pole of Plat =
34.6°N and Plon = 185.5°E with A95 = 6.5°, and the reversed polarity ChRM (D = 186.2°; I = 24.9°;
α95 = 3.5°; k = 31.1; N = 56) yields a paleomagnetic pole of Plat = 16.2°N and Plon = 194.8°E with
A95 = 5.0°. The data positions Baltica on equatorial latitudes, supporting the NENA (North Europe–
North America) connection between Baltica and Laurentia at 1.59–1.58 Ga. Furthermore, it is
demonstrated that NENA was valid also at 1.75, 1.46, and 1.26 Ga, forming the core of
Mesoproterozoic Nuna supercontinent.
4.3. Article III
Klein, R., Salminen, J., Mertanen, S., 2015. Baltica during the Ediacaran and Cambrian: A
paleomagnetic study of Hailuoto sediments in Finland. Precambrian Research 267, 94-105.
This article presents a new Late Neoproterozoic paleomagnetic pole for Baltica from a 272 m deep,
fully oriented sedimentary drill core in Hailuoto, Western Finland. The aim of the study was to obtain
a new paleomagnetic pole for Baltica, to test the Late Neoproterozoic paleogeographic positions of
Baltica, and to better understand the environmental conditions during that time. With a new Late
Neoproterozoic pole, a further aim is to explore if the large sways in the Ediacaran APWP can be
explained by high plate velocities alone.
Hailuoto Island is situated off the coast of the city of Oulu, Northern Finland (65°N, 25°E). The island
arose from the Bothnian Bay during the Quaternary post-glacial land uplift (Tynni and Donner, 1980).
The sedimentary bedrock of Hailuoto is covered by a Quaternary layer of loose sand of up to ca. 70
m thick. The bedrock consists of interbedded conglomerate, sandstone and mudstone up to a
maximum depth of ca. 560 m. The dominant rock type of the Hailuoto formation is a medium-grained,
pale pink or light greenish subarkose (Veltheim, 1969; Tynni and Donner, 1980). The depositional
age of the Hailuoto sediments is poorly constrained at 570-600 Ma based on fossil records (Tynni
and Donner, 1980).
30
Three components of magnetization were isolated from the Hailuoto sedimentary rocks with thermal
and alternating field (AF) demagnetization treatments. The ChRM component is a high
coercivity/unblocking temperature dual polarity component that passes a reversal test. The combined
observed ChRM component of the Hailuoto sediments (D = 334.2°; I = 44.4°; α95 = 7.2°; k = 16.5)
yields a paleomagnetic pole of Plat = 48.7°N and Plon = 241.1°E with A95 = 8.1°. The inclination
corrected direction (using f = 0.6 as a correction factor) of D = 334.4°; I = 57.7°; α95 = 5.8°; k = 25.2
yields a paleomagnetic pole of Plat = 60.5°N and Plon = 247.9°E with A95 = 7.6°. As it is a dualpolarity ChRM pole carried by both magnetite and hematite, with no resemblance to younger poles,
we interpret it as a primary component. A paleolatitude for Hailuoto of 38.3° was calculated from the
ChRM. Two secondary components were identified. The first is a low coercivity/blocking
temperature component with a remanent magnetization of D = 239.0°; I = 67.3°; α95 = 8.7° (N = 13
samples), which we interpret as drilling-induced remanent magnetization (DIRM).
The other
secondary component has a remanent magnetization of D = 49.4°; I = 34.9°; α95 = 8.6° (N = 5
samples) and is commonly seen in Fennoscandian formations.
The ChRM Hailuoto pole adds to the scattered Ediacaran paleomagnetic data of Baltica and indicate
large distances between other late Neoproterozoic and early Cambrian paleomagnetic poles, which
can be explained by high plate velocity. However, considering an older age for Hailuoto sediments,
the TPW and a non-actualistic geodynamo hypotheses cannot be ruled out. Reconstructions of Baltica
and Laurentia between 616 and 550 Ma is shown, which move Baltica from high latitudes (615 Ma),
over the polar region, to low latitudes (550 Ma), and Laurentia from low latitudes (615 Ma) to a polar
position (570 Ma) and back to an equatorial position (550 Ma). A low to mid latitude position of
Baltica determined by the Hailuoto paleomagnetic pole is in agreement with the earlier studies of
Hailuoto sediments which indicate a lack of glaciogenic sediments and a warm deposition
environment.
4.4. Article IV
Klein, R., Pesonen, L.J., Mänttäri, I., (submitted). A new late Paleoproterozoic pole for Baltica: A
paleomagnetic and geochronological study of the Keuruu diabase dykes, Central Finland.
Precambrian Research.
In this article a new paleomagnetic pole and isotope age for the late Paleoproterozoic (Svecofennian)
diabase dykes from Keuruu, Central Finland is presented. The motivation for this study is to obtain
late Paleoproterozoic key paleomagnetic data for Fennoscandia with precise U-Pb ages, to refine the
31
late Paleoproterozoic APWP of Baltica, and to present a paleogeographic reconstruction for the late
Paleoproterozoic that shows the onset of Nuna.
The Keuruu diabase dyke swarm is situated within the Central Finland Granitoid Complex (CFGC),
which was formed 1.89-1.86 Ga ago during the peak phase of the Svecofennian orogeny. Granitoids
are predominant rocks in the CFGC together with granodiorites, diorites and schists, as well as
numerous gabbroic bodies. These rocks are cut by the Keuruu diabase dyke swarm. An earlier
preliminary study of the dykes yielded a U-Pb (zircon) age of 1876±9 Ma, however the upper
intercept isochron age is based on only two zircon fractions.
The paleomagnetic results reveal a dual-polarity ChRM with a reversal asymmetry, which is
explained by an unremoved secondary component contaminating both N and R vectors. The normal
polarity ChRM (D = 346.1°; I = 37.0°; α95 = 4.8°; k = 57.2; N = 17) yields a paleomagnetic pole of
Plat = 47.9°N and Plon = 224.2°E with A95 = 4.0°, and the reversed polarity ChRM (D = 144.6.1°; I
= -23.3°; α95 = 13.4°; k = 33.5; N = 5) yields a virtual geomagnetic pole (VGP) of Plat = 34.0°N and
Plon = 247.9°E with A95 = 13.0°. The primary nature of magnetization in reversed polarity dykes is
supported by a baked contact test. As the normal polarity dykes and the unbaked host rocks show
similar paleomagnetic directions (similar ages), the baked contact test is illustrated by means of
remanent magnetization intensity decay behavior. A dyke showing reversed polarity is dated 1870±9
Ma by U-Pb. The youngest zircons from another dyke of normal polarity show a Pb-Pb age of 1868±7
Ma. A combined N and R pole of Plat = 45.4°N and Plon = 230.9°E with A95 = 5.5° positions
Fennoscandia at low latitudes (19°), and aligns with Laurentia in a way that would proceed into the
subsequent long lasting 1.8-1.2 Ga NENA configuration.
4.5. Author’s contribution
Article I: R. Klein was the leading author. He collected and prepared 10% of the samples and
measured 40% of the samples; the remaining samples had been collected, prepared and measured by
the second author, L.J. Pesonen with the help of Dirk Kettrup. Klein was responsible for the data
analysis, including the tilt test and inclination correction. Klein wrote most of the manuscript and
handled the review and proof stage.
Article II: R. Klein was the second author. He participated in the collection and preparation of the
samples in 2013, together with J. Salminen and S. Mertanen, as well as the paleomagnetic and rock
measurement of the samples. Klein was responsible for analyzing samples that had been collected by
L.J. Pesonen prior to the 2013 collection trip. Salminen was the leading author of the article, writing
32
most of the manuscript and handling the review and proof stage. Klein interpreted the paleomagnetic
results of the Rapakivi granites, participated in image preparation and participated in the proof
reading.
Article III: R. Klein was the leading author. He conducted most of the measurements and was
responsible for data analysis and interpretation. Klein did most of the writing, with significant input
by the co-authors. He was responsible for handling the manuscript during the review and proof stage.
Article IV: R. Klein was the leading author. Sample collection and preparation, and paleomagnetic
measurements had been carried out by L.J Pesonen with the help of R. Puranen and S. Raiskila.
Paleomagnetic data analysis and interpretation were carried out by Klein. Geochronological sample
preparation and measurements were carried out by Klein and I. Mänttäri, and geochronological data
analysis were done by Mänttäri. Klein wrote most of the manuscript.
5. Results and discussion
5.1. Paleomagnetic results obtained in this work
5.1.1. Diabase dykes
A dual polarity ChRM was obtained from Keuruu diabase dykes with a combined ChRM of D =
323.8°, I = 37.5°, k = 31.0, α95 = 5.8°. However the normal (D = 346.1°, I = 37.0°, k = 57.2, α95 =
4.8°) and reversed (D = 144.6°, I = -23.3°, k = 33.5, α95 = 13.4°) observed ChRM directions do not
pass the McFadden and McElhinny (1990) reversal test. The main possible cause for the asymmetry
between the N and R polarity Keuruu diabase dykes is interpreted to be an unremoved secondary
component. Furthermore, as the reversed direction is carried by only 5 dykes (compared to 17 dykes
of N polarity), the effect of unaveraged secular variation may therefore also contribute to the
asymmetry between normal and reversed distributions.
A reversed polarity dyke yielded a U-Pb zircon age of 1870±9 Ma and a normal polarity dyke yielded
a Pb-Pb age of 1868±7 Ma. The primary magnetization of the reversed component is supported by a
positive baked contact test. As the granodiorite host rock is only slightly older (1.89 Ga) than the
diabase dykes, and shows a similar normal polarity remanence direction as the normal polarity dykes,
it is not possible to distinguish between baked and unbaked host rock. However, baked and unbaked
host rocks can be distinguished based on the shapes of the remanent magnetization intensity decay
curves (Article IV), which support the primary nature of the N dyke direction as well.
33
Paleomagnetic results from Subjotnian intrusives of the Åland archipelago also show a dual polarity
ChRM with a combined ChRM of D = 14.9°, I = -2.6°, k = 14.0, α95 = 4.0°. The normal (D = 18.4°,
I = 19.9°, k = 13.0, α95 = 6.9°) and reversed D = 192.5°, I = 17.7°, k = 31.0, α95 = 3.5°) ChRM also
show a pronounced reversal asymmetry occurring in both quartz porphyry and diabase dykes. The
asymmetry is also explained by an unremoved secondary component. The main magmatic activity is
defined by U-Pb (zircon) data as 1580-1570 Ma, which is also regarded as the age of remanence,
supported by a positive baked contact test for both normal and reversed polarity dykes.
Non-antiparallel polarities
One of the reliability criteria for a paleomagnetic pole is the presence of reversals (Van der Voo,
1990), i.e. normal and reversed poles that are antiparallel in 95% confidence level. Both normal and
reversed directions of both Satakunta sandstone and Hailuoto sediments are antiparallel, and therefore
pass the reversal test.
Keuruu and Åland poles, however, show reversal asymmetry, i.e. their normal and reversed directions
are non-antiparallel. In both cases, it is interpreted to be a result of an unremoved secondary
component. This is apparent from the demagnetization data of several specimens. It was further tested
through vector addition, adding the secondary components to the expected original N and R ChRM’s
of Keuruu and Åland dykes. The respective combined Fisher means of the observed N and R ChRM’s
of Keuruu and Åland dykes (Table 2) were used as the expected original ChRM’s (Donadini, 2007).
In the case of Keuruu dykes, an R to N intensity ratio of 2:3 based on the mean R and N susceptibilities
(Table 2) was used for the expected original NRM values, assuming that susceptibility would reflect
the original NRM intensity ratio of the dykes prior to the partial secondary overprint. Since the
observed secondary direction is likely to be contaminated by PEF, we used the direction (D = 47°; I
= 64°; Pesonen et al., 2012a) obtained by great circle analysis. A secondary component intensity of
23% of the primary intensity resulted in an N direction of D = 341°; I = 39°, and an R direction of D
= 147°; I = -20°, which is within the 95% confidence of observed N and R directions (Table 2), with
a similar declination and inclination asymmetry.
The same procedure was followed for the Åland dykes, using an N to R intensiy ratio of 1:6, also
based on susceptibility values. The Fisher mean for the Åland secondary component (D = 19°; I =
57°) was used, with an intensity of 32% of the primary intensity. This resulted in an N direction of D
= 9°; I = 9°, and an R direction of D = 184°; I = 24°, which is within the 95% confidence of the
observed N and R directions (Table 2).
34
Table 2
Summary of ChRM results obtained from Articles I-IV.
Age
(Ma)
d.m.
B/N/n
N
1868±7
g(z)
17*/53/98
348.1
40.0
4.8
57
47.9
224.2
4.0
R
1870±9
g(z)
5*/18/46
146.7
-24.9
13.4
34
34.0
247.9
13.0
C
1870
21*/71/144
342.8
37.5
5.8
31
45.4
230.9
5.5
N
1569±3
g(z)
31*/150/272
18.4
19.9
6.9
13
34.6
185.5
6.5
R
1576±3
g(z)
56*/216/331
192.5
17.7
3.5
31
16.2
194.8
5.0
C
15801570
87*/366/603
14.9
-2.6
4.0
14
23.7
191.4
2.8
28.8
17.1
10.2
26
31.0
174.0
8.3
28.8
9.2
9.9
28
27.0
175.1
8.1
28.6
11.1
12.0
20
28.0
175.0
8.0
217.3
0.3
44.5
9
20.2
167.5
24.6
218.5
-17.0
19.2
42
28.3
163.4
14.9
Inc
218.8
-22.4
23.0
31
31.0
162.0
17.0
Obs
29.8
12.7
12.0
17
28.5
173.5
8.7
30.2
9.5
8.0
37
26.8
173.5
7,4
30.2
11.5
9.1
29
27.8
173.2
6.5
326.4
39.8
14.5
19
43.0
252.0
15.4
326.0
54.0
12.0
26
53.9
259.2
14.8
160.6
-43.8
8.5
17
49.5
235.1
9.7
161.0
-57.4
6.7
26
61.6
240.4
9
336.2
44.0
7.2
17
48.7
241.1
8.1
336.1
57.6
5.8
25
60.5
247.9
7.6
Rock unit
Pol
Keuruu diabase
dykes
Åland diabase and
quartz porphyry
dykes
Dr
(°)
Obs
Tilt
N
9*/30/72
Inc
Obs
Satakunta
sandstone
Tilt
Tilt
R
16301450
m,a
C
3*/7/16
10*/37/88
Inc
Obs
Inc
Hailuoto
sedimentary
drill core
Obs
Inc
Obs
Inc
N
R
C
7*/10
600570
F
19*/21
26*/31
Ir
(°)
α95
(°)
k
Plat
(°N)
Plon
(°E)
A95
(°)
Obs, Tilt, and Inc indicate observed, tilt-corrected and inclination corrected results from sedimentary formations. Pol =
polarity. d.m. shows dating method: g(z) = U-Pb (zircon), m = APWP, a = geology (cross-cutting), and f = fossils. B/N/n
= sites/samples/specimens (* marks level of statistics). Dr and Ir are declination and inclination of the remanent
magnetization component calculated for a common reference locality at Kajaani (64.1°N, 27.7°E). α95 = radius of the
95% confidence circle of the mean direction, k = precision parameter (Fisher, 1953). Plat and Plon = paleolatitude,
paleolongitude of the pole. A95 = 95% confidence circle of the paleopole.
35
5.1.2. Sedimentary rocks
Samples from the Satakunta sandstone formation, SW Finland) provided a stable ChRM component
of D = 27.8°, I = 3.9°, k = 28.9, a95 = 9.1° after tilt- and inclination- corrections. Results were obtained
from 10 outcrops of which 7 outcrops yielded normal polarity directions, 2 outcrops yielded both
normal and reversed polarity directions, and 1 outcrop yielded reversed polarity directions only. The
normal and reversed directions pass the McFadden and McElhinny (1990) reversal test. In addition,
the results passes a fold test as well as an inverse baked contact test. The sandstone is partially baked
by 1264±12 Ma (Suominen et al., 1991) olivine diabase dykes and sills. An earlier paleomagnetic
study on Satakunta sandstone (Neuvonen, 1973) was unable to isolate the sandstone component, and
interpreted the secondary diabase overprint as the ChRM of the sandstone. As a result the
paleomagnetic age of the Satakunta formation was estimated at 1400-1300 Ma. The paleomagnetic
age based on the new ChRM is estimated at 1630-1450 by comparing the new Satakunta pole to
reliable, dated poles (Fig. 3).
A dual polarity ChRM was obtained from a 272 m deep, fully oriented sedimentary drill core in
Hailuoto, Western Finland. The depositional age of the Hailuoto sediments is poorly constrained at
600-570 Ma based on fossil records. The normal and reversed polarity directions pass the McFadden
and McElhinny (1990) reversal test. The combined observed ChRM component is D = 334.2°; I =
44.4°; k = 16.5; α95 = 7.2° and the inclination corrected direction (f = 0.6) is D = 334.4°; I = 57.7°; k
= 25.2; α95 = 5.8°. There was no possibility for field tests. However, as it is a dual-polarity ChRM
carried by both magnetite and hematite, with no resemblance to younger events, we interpret it as a
primary component.
5.1.3. Secondary component
Both Satakunta sandstone and Hailuoto sediments show a secondary component with intermediate
downward inclinations and northeasterly declinations. In the case of Satakunta sandstone it is D =
27.2° I = 33.7° α95 = 13.0° and is observed in addition to the Postjotnian component mentioned above
(section 4.1.). In the case of the Hailuoto sedimentary formation the secondary component is D =
49.4° I = 34.9° α95 = 8.6°. Secondary components are also observed in Keuruu and Åland dykes,
however it is difficult to isolate the present Earth field component (PEF) and secondary component,
as they co-exist in a similar coercivity range of ≤ 20 mT or a blocking temperature range of ≤ 300°C.
The secondary component, however, is also carried at coercivities of >200 mT and temperatures of >
500°C in some of the samples, where they co-exist with, and contaminate the ChRM. The Fisher
36
mean of the secondary component at low coercivity/blocking temperature and at high
coercivity/blocking for both Keuruu and Åland cases are shown in Table 3. A great circle analysis on
Keuruu dykes (Pesonen et al., 2012a) isolated a secondary component of D ≈ 47°, I ≈ 64°. Raiskila
et al. (2011) obtained a similar direction (D=42°, I = 64°) from the 1.14 Ga Keurusselkä impact
structure ca. 20 km South of Keuruu, and interpreted it as an impact component. This may be a
possible source of the secondary component in Keuruu dykes, however there is no significant proof
that this is an impact component. Similarly, the ca. 0.55-0.45 Ga Lumparn impact (Tynni, 1982)
cannot be ruled out as a possible source of the secondary component in the Åland case.
An intermediate to steep NE component is widely observed in Baltica in ealier paleomagnetic studies
(e.g. Mertanen et al., 1989; Preeden et al., 2009; Salminen et al., 2014). The paleomagnetic pole
derived from this component corresponds with Permian to Triassic (ca. 300-200 Ma) poles of Europe
(Torsvik et al., 2012). Preeden et al. (2009) interpret the cause of this component to be the migration
of orogenic fluids related to the Hercynian orogeny at the southern margin of Baltica and/or the
Uralian orogeny at the eastern margin.
Table 3
Fisher means of secondary components obtained in this work
B/N/n
D
(°)
I
(°)
k
Plat
(°N)
Plon
(°E)
A95
(°)
<20mT
<300°C
19*/47/71
24.8
68.4
8.0
18.5
75.7
131.1
11.5
>30mT
>500°C
18*/42/57
9.1
51.6
8.7
16.6
61.9
188.1
8.5
<20mT
<300°C
24*/52/52
38.8
68.7
6.3
22.9
67.6
110.2
12.0
>30mT
>500°C
10*/20/20
19.0
57.0
8.5
33.6
65.1
164.7
19.7
Component B
8*/29/44
27.7
33.7
13.0
19.2
44.3
163.7
11.4
Baked component
13*/50/112
34.0
-40.7
8.1
27.5
-1.1
350.6
12.3
5*/7
49.4
34.9
8.6
80.7
34.1
144.7
6.9
Rock unit
α95
(°)
Keuruu diabase dykes
Åland diabase and quartz
porphyry dykes
Satakunta sandstone
Hailuoto sedimentary drill core
B/N/n = sites/samples/specimens (* marks level of statistics). Dr and Ir are declination and inclination of the remanent
magnetization component. α95 = radius of the 95% confidence circle of the mean direction, k = precision parameter
(Fisher, 1953). Plat and Plon = paleolatitude, paleolongitude of the pole. A95 = 95% confidence circle of the paleopole.
37
5.2.Aparent polar wander paths for Baltica (1.87-0.57 Ga)
The ChRM poles of Keuruu diabase dykes, Åland intrusives, and Satakunta sandstone are plotted
with selected Paleo- and Mesoproterozoic poles (Fig. 3). The poles used are classified into highly
reliable A grade poles (well determined age and positive field stability test) and seemingly reliable B
grade poles (well determined age or positive field test). Questionable to unreliable C and D grade
poles (neither well determined age nor positive field test) are not included here. In Fig. 3 the late
Paleoproterozoic to Mesoproterozoic APWP of Baltica starts at an overlapping cluster of poles
consisting of the Tsuomasavarri (1931 Ma), mean 1880 Ma Svecofennian, Keuruu dykes (1870 Ma),
mean 1800 Ma Svecofennian, Hoting gabbro (1786 Ma), Shoksha sandstone (1770 Ma), and
Ropruchey sill (1751 Ma) poles; all positioned at intermediate latitudes. The APWP then moves down
to the subjotnian (ca. 1.65-1.45 Ga), Mashak (1376 Ma), and postjotnian (1.26 Ga) poles at low to
equatorial latitudes, anchored by Keuruu dykes (1870 Ma), Åland intrusives (1578 Ma), Satakunta
dykes (1578 Ma), Lake Ladoga (1452 Ma), and 1258 Ma mean postjotnian key poles.
Fig. 3. Late Paleoproterozoic to late Mesoproterozoic APWP for Baltica. A grade poles are marked in red, B grade poles
are marked in orange, and C grade poles marked in green. The poles of this work are shown with arrows.
38
The late Mesoproterozoic and Neoproterozoic is an enigmatic period for Baltica and contains few
reliable paleomagnetic poles and has large age gaps between them, making a continuous APWP
between the Precambrian and Paleozoic impossible. In this dissertation, an APWP of only the
transition from Precambrian to Ordovician (615 – 475 Ma) is discussed. The ChRM pole of Hailuoto
is plotted with selected A and B grade Neoproterozoic – Cambrian poles (Fig. 4). Although the exact
deposition age of the Hailuoto sediments is unkown, for the purpose of this discussion the age range
proposed by Tynni and Donner (1980) of 600 - 570 Ma is used, which places the age of the Hailuoto
pole between the Egersund (616 Ma) and the Verkhotina (Popov et al. 2005) and Zolotitca (Iglesia
Llanos et al. 2005) poles (ca. 550 Ma). This forms a loop in the APWP that starts at the Egersund
pole (616 Ma), jumps to the Hailuoto pole (ca. 600 - 570 Ma), and follows the poles of Krivaya Luka,
Kurgashlya and Bakeevo formations (570 – 560 Ma) (Lubnina et al. 2014), continues to the 550 Ma
Verkhotina (Popov et al. 2005) Zolotitca (Iglesia Llanos et al. 2005) and Winter Coast (Popov et al.,
2002) poles, and ends at the 478 Ma St. Petersburg limestone pole (Smethurst et al. 1998) close to
the Egersund pole (illustrated with black dotted line in Fig. 4).
Fig. 4. Late Neoproterozoic to Ordovician APWP for Baltica (left) and Laurentia (right). Neoproterozoic poles marked
in yellow, Cambrian poles marked in green, and Ordovician poles marked in orange.
The angle of the Baltica APWP segment between the Egersund dyke pole (616 Ma) and the Hailuoto
pole is 98° for the uncorrected pole and 86° for the inclination corrected pole (f=0.6), which infers a
plate motion of between ca. 24 and 21 cm/yr, respectively, considering the lower Hailuoto age limit
39
of 570 Ma. This is consistent with the upper speed limit for plate tectonics (between 20 and 25 cm/yr)
suggested by Meert et al. (1993) and Gurnis and Torsvik (1994). The APWP of Baltica between the
550 Ma mean pole and the St Peterburg (478 Ma) (Smethurst et al. 1998) pole (Fig. 4) yields an 87°
angle which results in a velocity of ca. 16 cm/yr when considering only A and B grade poles.
Although the the large sways in the APWP of Baltica can be explained by high plate velocity alone,
considering the lower age 570 Ma for Hailuoto sediments, when considering an older age for
Hailuoto, high plate velocity becomes unrealistically high (ca. 65 cm/yr at 600 Ma), and other
explanations, such as TPW (Mitchell et al., 2011) and non-actualistic geodynamo (Abrajevitch and
Van der Voo, 2010) hypotheses should be considered.
True polar wander (TPW) (Evans, 2003; Mitchell et al., 2011) and a non-actualistic geodynamo
(Abrajevich and Van der Voo, 2010) are two hypotheses given as possible mechanisms for the rapid
continental motion during the Ediacaran. True polar wander is the rotation of the Earth with respect
to its spin axis causing the geographic locations of the North and South poles to change or “wander”.
Such a shift in poles results in a systematic change in the APWP’s of all cratons.
In the Laurentia APWP (Fig 4.), the segment between the 613 Ma Long range pole (Murthy et al.,
1992) and the 585 Ma Grenville B+E pole (Halls et al., 2015) form an 88° angle. This is similar to
the segment between the 616 Ma Egersund pole (Walderhaug et al., 2007; Bingen et al., 1998) and
the Hailuoto pole in the Baltica APWP, and therefore supports the TPW hypothesis. However, the
Laurentia APWP is inconclusive, as there are other high quality poles for ca. 585 Ma at different
positions that would result in smaller APWP angles with the Long Range pole. Furthermore, the
angles between poles younger than 585 Ma differ between Baltica and Laurentia, and therefore
negates TPW as the main cause of large APWP segments after 585 Ma. However, in the case of
Laurentia, the large angle of 66q of the APWP segment between Catoctin Basalts (572 Ma) and SeptIles intrusions (564 Ma), which equates to 105 cm/yr, cannot be explained by high plate velocity
alone.
The non-actualistic geodynamo hypothesis by Abrajevitch and Van der Voo (2010), which postulates
an alternation of the geomagnetic dipole axis between a co-axial and an equatorial alignment, partly
relies on problematic poles for both Laurentia and Baltica. The Grenville dykes yield both a high and
low latitude pole (Halls et al., 2015; Murthy, 1971). Halls et al. (2015) propose that the high latitude
pole is the result of an intermediary stage in the reversal of an axial dipole field. In the case of Baltica
the model is based on the Alnö carbonate complex (584 r 7 Ma) (Meert et al., 2007; Piper, 1981) in
which both a high and a low latitude component co-exist in the same unit. However the primary nature
40
has not been proven for either. Furthermore, the Fen carbonate complex pole (Meert et al., 1998),
another pole in the model, lacks stability tests and is likely a Permo-Triassic overprint (Meert, 2014b).
5.3. Paleogeographic reconstructions (1.87-0.57 Ga)
Mesoproterozoic
Paleogeographic reconstructions at 1.87 Ga and 1.57 Ga are presented here based on the Keuruu
diabase dyke and Åland intrusives paleomagnetic poles, respectively, with other poles listed in Table
4. The 1.87 Ga reconstruction presented here (Fig. 5) indicates the onset of the supercontinent Nuna,
which was assembled by at least ca. 1650 – 1580 Ma (Pisarevsky et al, 2014a).
At 1.87 Ga, Baltica and Laurentia were not yet assembled. Sarmatia and Volgo-Uralia collided with
the Fennoscandian Shield only between 1.82 and 1.80 Ga along the Russian collisional belt to form
Baltica (Bogdanova et al., 2008, 2015; Elming et al., 2010). The Trans-Hudson orogeny, connecting
Superior-Nain and Slave-Rae-Hearne blocks of Laurentia, occurred between 1.91 and 1.81 Ga
(Hoffman, 1988). For ease of illustration, Laurentia is shown as a united craton at 1.87 Ga (Fig. 5),
however the position of Superior and Slave poles do not support a united craton at this time yet. In
the 1.87 Ga reconstruction (Fig. 5), Fennoscandia is positioned at low latitudes and aligns with
Laurentia in a way that would proceed into the 1.8-1.2 Ga NENA configuration, which is established
in the 1.57 Ga reconstruction (Fig. 6).
Although there are no high quality paleomagnetic poles for Amazonia at ca. 1.87 Ga and ca. 1.57 Ga,
interpolations were used in both cases (Table 4), positioning Amazonia at near equatorial positions.
The position of Amazonia in the 1.87 Ga reconstruction (Fig. 5) resembles a SAMBA-type
configuration (South America – Baltica), in which northern Amazonia is attached to SW Baltica
(Johansson, 2009). The SAMBA configuration is supported by the 1.79 Ga Avanavero intrusions pole
(Bispo-Santos et al., 2014b; Pisarevsky et al., 2014a). The position of Amazonia at 1.57 Ga (Fig. 6)
is supported by the 1789 Ma Colider Volcanics pole (Bispo-Santos et al., 2008) and the 1439 Ma
Aquapei pole (Elming et al., 2009; Geraldes et al., 2014), both of which overlap with the Åland pole
of Baltica.
Siberia possibly formed part of the Nuna core (Wingate et al., 2009; Evans and Mitchell, 2011). The
1863 Ma Upper Akitkanchaya pole indicates a low latitudinal position of Siberia with a large distance
between Siberia and Laurentia, similar to the position of Siberia at ca. 1770 - 1720 Ma (Pisarevsky
et al., 2014a). In the 1.57 Ga reconstruction (Fig.6), a direct connection between Siberia and Laurentia
41
is shown as in the reconstruction by Mitchell and Evans (2011). Rainbird et al. (1998) identified
northwest Canada as the source region for Mesoproterozoic zircons in mid-late Riphean sandstone in
southeastern Siberia, supporting a close connection between Laurentia and Siberia during the Mesoand Neoproterozoic.
Fig. 5A. Paleogeographic reconstruction of continents at 1.87 Ga based on results from Article IV. 1.88 Ga
reconstructions by Pesonen et al. (2012) and Belica et al. (2014) added for comparison. FSc = Fennoscandia, La =
Laurentia, Am = Amazonia, Si = Siberia, WAU = Western Australia, In = India, NC = North China, Ka = Kalahari. The
colour of the poles correspond with the colour of craton outlines. Closed and open Laurentia poles indicate Superior and
Slave poles, respectively. B. Postion of Keuruu pole and selected poles from Laurentia, Siberia and Amazonia (Table 4)
in the 1.88 Ga paleogeographic reconstruction by Pesonen et al. (2012). C. Paleogeographic reconstruction modified after
Belica et al. (2014) with alternative position for Siberia (Baltica and Laurentia the same as in B).
The 1.57 Ma Nuna reconstruction (Fig. 6) shows the North Australian craton located at NW Laurentia
in a SWEAT-like configuration (e.g. Pisarevsky, 2014a; Furlanetto et al., 2013; Zhang et al., 2012;
Betts et al., 2008, 2009). The North Australian craton is currently the only Australian craton with
reliable paleomagnetic results for the 1.57 Ma reconstruction, however the North Australian, West
42
Australian and South Australian cratons were either already assembled into present day Australia, or
were in close proximity to each other since ca. 1800 Ma (Pisarevsky et al., 2014a and references
therein; Li and Evans, 2011). The Western Australian craton is situated at low latitudes in the 1870
Ma reconstruction (Fig. 5) from where it possibly moved northwards to form the SWEAT-like
configuration with Laurentia during the Mesoproterozoic.
Fig. 6A. Paleogeographic reconstruction of continents at 1.57 Ga based on results from Article II. 1.58 Ga reconstructions
by Pisarevsky et al. (2014a) and 1.59 Ga reconstruction by Zhang et al. (2012) added for comparison. Ba = Baltica, La =
Laurentia, Am = Amazonia, Si = Siberia, NAU = Northern Australia, In = India, NC = North China, Ka = Kalahari, C =
Congo, SF = São Francisco. The colour of the poles correspond with the colour of craton outlines. Baltica rotated to
Laurentia: Elat = 47.5°, Elon = 1.5°, angle = 49° (Evans and Pisarevsky, 2008). Siberia rotated to Laurentia: Elat = 76.8°,
Elon = 106.2°, angle = 146.6° (Evans and Mitchell, 2011). North Australia rotated to Laurentia: Elat = 31.5°, Elon = 98°,
angle = 102.5° (Evans and Mitchell, 2011). India is rotated to the absolute framework: Elat = -13.21°, Elon = 136.86°,
angle = 197.48° (Pisarevsky et al., 2014a). B. Postion of Åland pole and selected poles from Laurentia, Siberia and
Amazonia (Table 4) in the 1.58 Ga paleogeographic reconstruction by Pisarevsky et al. (2014a). C. 1.59 Ga reconstruction
by Zhang et al., 2012) with same euler rotations for Baltica, Laurentia, Amazonia, and Siberia as in A.
43
Both the 1.88 Ga Dharwar dykes pole and the 1.47 Ga Lakhna dykes pole position India at an
equatorial position. In the 1.87 Ga reconstruction (Fig. 5) India is positioned south of Fennoscandia
based on geological matches between India and Sarmatia (e.g. Pisarevsky et al., 2013). The low
latitudinal position of North China during the late Paleoproterozoic and Mesoproterozoic is supported
by paleomagnetic data (e.g. Table 4) and ca. 1.7 Ga stromatolites (Zhang, 1988). In the 1.87 Ga
reconstruction (Fig. 5) North China is positioned close to Western Australia, as proposed by Zhang
et al. (2012).
Based on the 1.77 - 1.27 reconstructions of Pesonen et al. (2003), Jacobs et al. (2008), and Pisarevsky
et al. (2014a) in which Kalahari is surrounded by ocean, Kalahari is positioned as a “lone” continent
in the Southern hemisphere in the 1.87 Ga reconstruction (Fig. 5). Based on a mismatch of the 1.87
– 1.88 Ga Mashonaland pole from the Zimbabwe craton and the 1.88 Ga Post-Waterberg pole from
the Kaapvaal craton (Table 4), Hanson et al. (2011) proposed that Kalahari did not assemble until
after 1.87 Ga. However the qualtity of the Post-Waterberg pole is questionable as it is not supported
by a field test nor a reversal test. We therefore keep Kalahari as a unified craton in the 1.87 Ga
reconstruction.
Table 4
Paleomagnetic poles used in the reconstructions
Plat
(°N)
Plon
(°E)
A95 (°)
Age
(Ma)
Q(1-7)
Reference
Keuruu diabase dykes
45
231
5.5
1870
1101110=5
Article IV
Åland intrusives
24
191
2.8
15801570
1111111=7
Article II
Mean Postjotnian intr.
-1.8
159
3.4
1258
1111101=6
Pesonen et al., 2003
Laanila-Ristijarvi dykes
-2
212
13.8
1044
0011100=3
Mertanen et al., 1996
Egersund dykes
-31
224
15.7
616 ± 3
1111101=6
Walderhaug et al. (2007),
Bingen et al. (1998)
obs.
-49
61
8.1
600-570
0110111=5
Article III
i.c.
-61
68
7.6
Formation
Baltica
Hailuoto
sediments
44
Mean 570 Ma
-52
118
570
1111111=7
1
obs.
-30
111
550
1111111=7
2
i.c.
-42
116
Molson dykes B+C2
28.9
218
3.8
1877 +7/- 1111111=7
4,
1884 ± 2
Halls and Heaman (2000)
Zhai et al. (1994)
Halls and Evans (2010)
MEAN 1:
Haig/Flaherty/Sutton
1
246
3.9
1870 ± 1
1111111=7
Luleå working group (2009)
MEAN 2:
Seton/Akaitcho/Mara
-6
260
4.0
1885 ± 5
1111111=7
Mitchell et al. (2010)
MEAN 3:
-12
Kahochella/Peacock Hills
285
7.0
1882 ± 4
1100111=5
Mitchell et al. (2010)
MEAN 4: Pearson A
/Peninsular sill/
Kilohigok basin sill
-22
269
6.0
1870 ± 4
1101111=6
Mitchell et al. (2010)
Stark formation
-15
215
4.9
1876±10
1110111=6
Bingham and Evans (1976)
Western Channel diabase
9.0
245.0
6.6
1590± 3
1101101=5
Irving et al. (1972)
Hamilton and Buchan (2010)
MacKenzie dykes
4
190
5.0
1267
1111101=6
Buchan et al. (2000)
Freda sandstone
2
179
4.2
1050
1110100=4
Henry et al. (1977)
Long Range dykes
-19
175
17.4
615 ± 2
1111111=7
Murthy et al.(1992)
Grenville dykes (B+E)
-57
48
9.1
585
1111110=6
Halls et al. (2015)
Baie de Mouton complex
(A)
-43
153
12.0
583 ± 2
1110001=4
McCausland et al. (2011)
Baie de Mouton complex
(B)
34
142
15.4
Catoctin Basalts
-42
117
17.5
572 ± 5
1111111=7
Meert et al. (1994b)
Sept Iles intrusion
20
142
6.7
565 ± 4
1111100=5
Tanczyk et al. (1987)
Skinner Cove volcanics
16
157
9.0
550 ± 3
1111100=5
McCausland and Hodych (1998)
Surumu volcanics
27
235
8.1
1970 ± 10 1111101=6
Bispo-Santos et al., 2014a
Avanavero mafic rocks
48
208
9.2
1789 ± 3
Bispo-Santos et al., 2014b
Mean 550
Ma
Laurentia
1010010=3
Amazonia
45
1111101=6
Colider volcanics
63
119
9.5
1789 ± 7
1110111=6
Bispo-Santos et al., 2008
Aguapei sills/dykes
65
101
10.1
1439 ± 7
1110111=6
Elming et al., 2009
Geraldes et al., 2014
Nova Guarita dykes
48
66
6.6
1419 ± 4
1111111=7
Bispo-Santaos et al., 2012
Nova Floresta formation
25
165
5.5
1200 ± 7
1110101=5
Tohver et al. (2002)
Rio Negro basement
5.2
7.5
1035
1000011=3
Bettencourt et al. (1996)
Upper Akitkanchaya
igneous and sedimentary
rocks
-22
98
5.2
1863 ± 9
1110100=5
Didenko et al. (2009)
Sololi-Kyutingde
intrustions
-37
73
10.4
1473 ±
24
Linok formation
15.2
76
7.5
1050
0111111=6
Gallet et al. (2000)
Frere formation
45
40
1.8
1890±10
1110111=6
Williams et al. (2004)
Rasmussen et al. (2012)
Gnowangerup-Fraser
dykes
56
324
6.3
1215
1111111=7
Pisarevsky et al. (2014b)
Balbirini dolomite (upper -52
part)
176
7.5
1589± 3
1110110=5
Idnurm (2000)
Alcurra dykes and sills
3
80
8.8
1077
1111001=5
Schmidt et al. (2006)
Dharwar 1.88 Ga dykes
37
334
3.2
1885± 3
1111110=6
Belica et al. (2014)
Lakhna dykes
41
121
20.5
1466± 3
1110111=6
Pisarevsky et al. (2013)
Majhgawan Kimberlite
37
213
12.2
1074±14
1010100=3
Gregory et al. (2006)
Mashonaland dolerites
-8
158
5.0
1871.9 ± 1111111=7
2.2,
1881.1
+5.9/-5.1,
1882.7
+1.6/-1.5
Bates and Jones (1996)
Hanson et al. (2011)
Söderlund et al. (2010)
Post-Waterberg sills
-9
194
18.0
1875±4
Hanson et al. (2004)
Siberia
Wingate et al., (2009)
W. Australia
N. Australia
India
Kalahari
46
1110101=5
Port Edward pluton
-7
328
4.4
1025 ± 8
1110001=4
Gose et al. (2004)
Bahia dykes
8
6
11.9
1500
Late Kibaran intrusives
-17
113
7.7
1236
1110001=4
Meert et al. (1994a)
281
7
1078±18
0111101=5
D’Agrella-Filho et al. (2004)
Congo
Salminen et al. (submitted)
São Francisco
Salvador and Olivença
9
dykes (R-polarity results)
North China
Taihang dykes
36
246
3.0
1769±3
1111101=6
Halls et al. (2000)
Yangzhuang formation
17
214
5.7
1560
0111111=6
Wu et al. (2005)
Pei et al. (2006)
Yanliao mafic sills
-6
180
4.3
1323±7
1111101=6
Chen et al. (2013)
Zhang et al. (2012)
1060 Ma interpolation
21
217
ca. 1060
3
Plat, Plon is latitude and longitude of the paleomagnetic pole. A95 = 95% confidence circle of the pole, Q(1-7) = Van der
Voo reliability criteria of paleomagnetic poles
References:
1.The 570 Ma mean pole was calculated from the Krivaya Luka, Kurgashlya and Bakeevo poles (Lubnina et al., 2014).
2. The 550 Ma mean pole was calculated from the Verkhotina comp. Z (Popov et al., 2005) and Zolotitca comp. B
(Iglesia Llanos et al., 2005) poles.
3. The 1060 Ma North China pole was interpolated from the ca. 1175 Ma Tieling sills (Zhang et al., 1991) and ca. 950
Ma Qingbaikouan (Zhang et al., 2006) poles.
Reconstructions at 1.25 Ga and 1.05 Ga are also presented here (Fig. 7), although no new poles were
obtained in this study for this period. In both the 1.25 Ga and 1.05 Ga reconstructions Baltica and
Laurentia are connected; in the 1.25 Ga reconstruction Baltica and Laurentia is in the 1.78-1.27 Ga
NENA configuration, whereas in the 1.05 Ga reconstruction western Norway is attached to eastern
Greenland. Furthermore, Amazonia is close to Baltica in the 1.25 Ga reconstruction, but is rotated
towards Laurentia in the 1.05 Ga reconstruction. The rotational changes of Baltica and Amazonia
between 1.25 Ga and 1.05 Ga, possibly forming the Grenville collisional belt, mark the transition to
Rodinia (Yakubchuk, 2010; Ibanez-Mejia et al., 2011; Roberts, 2013).
47
Fig. 7 (left). Paleogeographic reconstruction of continents at 1.25 Ga. 1.27 Ga reconstructions by Pisarevsky et al. (2014a)
and Evans and Mitchell et al. (2011) added for comparison. Ba = Baltica, La = Laurentia, Am = Amazonia, Si = Siberia,
Au = Australia (WAU and NAU = Western and Northern Australia, respectively), In = India, NC = North China, C =
Congo, SF = São Francisco. Baltica rotated to Laurentia: Elat = 47.5°, Elon = 1.5°, angle = 49°(Evans and Pisarevsky,
2008). Siberia rotated to Laurentia: Elat = 76.8°, Elon = 106.2°, angle = 146.6° (Evans and Mitchell, 2011). Amazonia
rotated to Laurentia: Elat = 53.93°, Elon = -68.94°, angle = 122.19° (based on SAMBA) (right). Paleogeographic
reconstruction of continents at 1.05 Ga. 1.05 Ga reconstructions by Evans (2009) and Li et al. (2008) added for
comparison. Ba = Baltica, La = Laurentia, Am = Amazonia, Si = Siberia, Au = Australia (WAU and NAU = Western and
Northern Australia, respectively), In = India, NC = North China, C = Congo, SF = São Francisco, Ka = Kalahari.
From the 1.27 Ga McKenzie dykes pole (Buchan et al. 2000) Laurentia moves to mid to high latitudes,
based on ca. 1.1 Ga paleomagnetic poles (e.g. Ernst and Buchan, 1993; Swanson-Hysell et al., 2014),
and gradually back to low latitudes, where it is again connected with Baltica at 1.05 Ga. Baltica lacks
good quality poles for this time period. The 1.12 Ga Salla dyke VGP (Salminen and Pesonen, 2007)
does not support the NENA configuration, however it closely resembles the secondary component
(component B), discussed above, commonly observed in Fennoscandian formations, and may
possibly not present the actual position of Baltica during that time. However some poor quility poles
from Sweden (Bylund, 1992) and Ukraine (Mikailova, 1989) suggest an excursion of Baltica to mid
latitudes, similar to Laurentia.
48
No good quality paleomagnetic data are available for Siberia at ca.1.25 Ga. Based on the 1.47 Ga
Sololi-Kyutingde (Windgate, 2009), and 1.05 Ga Linok (Gallet et al. 2000) pole positions, a
connection between Siberia and Laurentia is plausible at 1.25 Ga. The 1.05 Ga Linok pole, however
positions Siberia at higher latitudes and does not support a southeast Siberia- northwest Laurentia
connection, which contradicts geological support by Rainbird (1998).
The Australian cratons are positioned in the northern hemisphere in the 1.25 Ga (not shown in Fig.
7A) and 1.05 Ga reconstructions, alongside western Laurentia. Li et al. (2008) proposed that the
Australian cratons existed in close proximity to each other but in a different configuration than today,
until ca. 650 Ma. This notion is followed in the 1.25 Ga and 1.05 Ga reconstructions, as a presentday configuration would imply an unusual vertical axis rotation between 1.25 Ga and 1.05 Ga.
Neoproterozoic
At ca. 615 Ma Baltica and Laurentia were possibly still connected with Baltica at intermediate
latitudes and Laurentia at low latitudes (Fig. 8). The 613 Ma Long Range dykes (Laurentia) and the
616 Ma Egersund dykes are rift-related (Bingen et al. 1998, Puffer 2002), and may mark the start of
the break-up between Baltica and Laurentia. The orientation of Baltica changes by ca. 180° between
615 Ma and 600-570 Ma (Article III), suggesting that Baltica moved across the polar region, which
is consistent with earlier reconstructions (Meert, 2014; Li et al., 2013; McCausland et al., 2011). In
this model we propose two possible positions for Laurentia at 585 Ma, based on two high quality
paleomagnetic poles (Fig. 8). Considering the Grenville B+E pole (Halls et al. 2015), Laurentia
followed Baltica over the polar region, suggesting that these two cratons were still connected at ca.
585 Ma. Based on the Baie de Moutan A pole, Baltica and Laurentia had already separated by ca.
585 Ma. The Baie de Moutan A poles also positions Laurentia at high latitudes, but with eastern
Laurentia facing Baltica. At ca. 570 Ma Baltica is positioned at intermediate to low latitudes where
it remains at 550Ma. Laurentia, however remains at the polar region at ca. 570 Ma, and at 550 Ma is
positioned at equatorial latitudes.
49
Fig. 8. Reconstructions of Baltica and Laurentia between 615 -550 Ma. Baltica shown observed (obs.) and inclination
corrected (i.c.) positions in 550 Ma reconstruction.
The polar position of Laurentia at 570 Ma (Fig. 8) and the high latitude model of early Gondwana at
575 Ma (Pisarevsky et al. 2008), as well as the low to intermediate latitude of Baltica presented in
this study and warm deposition environment of Hailuoto sediments (Tynni and Donner, 1980),
suggest that no global scale glaciation event, as presented in the snowball Earth model (Kirschvink,
1992), took place at that time. These conditions further indicate zonal surface temperature gradients
similar to today, i.e. warm equatorial areas and cold polar areas, unlike the high obliquity model
presented by Williams (1975, 1993, 2008), which proposes an obliquity of the Earth’s spin axis with
respect to the ecliptic of >54° resulting in warm polar areas and ice ages in the tropics.
6. Conclusions
New paleomagnetic results were obtained from samples from 1.87 Ga Keuruu diabase dykes (Central
Finland), 1.58-1.57 Ga Åland intrusives (Åland archipelago, SW Finland) and ca. 0.6-0.57 Ga
Hailuoto sedimentary formation (Western Finland). In addition to these results, new paleomagnetic
results were obtained from the Satakunta sandstone formation (SW Finland), providing a new relative
(paleomagnetic) age estimate of ca. 1.63 – 1.45 Ga, older than previously thought (ca. 1.4-1.3 Ga).
Results from Åland intrusives and Satakunta sandstone place Baltica at equatorial positions and
support the NENA configuration between Baltica and Laurentia during the Mesoproterozoic. Results
from the older Keuruu diabase dykes positions the Fennoscandian Shield at low latitudes and aligns
50
with Laurentia in a way that would naturally proceed into the 1.8-1.2 Ga NENA configuration. The
primary nature of the Åland and Keuruu ChRM is supported by U-Pb rock ages, and a positive baked
contact tests, yielding two new key poles to the late Paleoproterozoic to late Mesoproterozoic APWP.
Results from the Hailuoto sediments add to the scattered Ediacaran paleomagnetic data of Baltica and
indicate large distances between other late Neoproterozoic and early Cambrian paleomagnetic poles,
implying rapid movement of Baltica from high latitudes at 615 Ma, over the polar region, to low
latitudes at 550 Ma. A low to mid latitude position of Baltica determined by the Hailuoto
paleomagnetic pole is in agreement with earlier geological studies of Hailuoto sediments that show
the lack of glaciogenic sediments and indications of a warm depostion environment.
Paleogeographic reconstructions are presented for 1.87 Ga, 1.57 Ga, 1.25 Ga, and 1.05 Ga. The 1.87
Ga reconstruction do not support the NENA or SAMBA configurations, but indicates the onset of
Nuna. The 1.57 Ga and 1.25 Ga reconstructions indicate similarities in the relative positions between
Baltica, Laurentia, Amazonia, Siberia. The transition between Nuna and Rodinia is marked by a
rotation between Laurentia, and Baltica and Amazonia (and possibly West Africa), forming the
Grenville collisional belt.
In all these cases a secondary component with shallow to intermediate inclinations and northeasterly
declinations were observed. It has been observed widely by previous authors and has been interpreted
as a Phanerozoic overprint. In the case of Keuruu and Åland, meteorite impact (from Keurusselkä
and Lumparn impacts, respectively) cannot be ruled out as a possible cause of the secondary
component. The secondary component is mostly carried by low coercivities/blocking temperatures,
but sometimes also at high coercivities/blocking temperatures causing asymmetry in the dual polarity
results of Keuruu and Åland diabase dykes.
51
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