UNIVERSITY OF HELSINKI
DEPARTMENT OF PHYSICS
REPORT SERIES IN GEOPHYSICS
No 59
PALEOMAGNETIC AND ROCK MAGNETIC STUDY WITH
EMPHASIS ON THE PRECAMBRIAN INTRUSIONS AND IMPACT
STRUCTURES IN FENNOSCANDIA AND SOUTH AFRICA
Johanna Salminen
HELSINKI 2009
UNIVERSITY OF HELSINKI
DEPARTMENT OF PHYSICS
REPORT SERIES IN GEOPHYSICS
No 59
PALEOMAGNETIC AND ROCK MAGNETIC STUDY WITH
EMPHASIS ON THE PRECAMBRIAN INTRUSIONS AND IMPACT
STRUCTURES IN FENNOSCANDIA AND SOUTH AFRICA
Cover picture:
Top left: Illustration of Earth’s magnetic field
Top right: Chinese compass (Si Nan) pointing to the South, used already 200 BA.
Bottom left: Salla dyke, NE Finland
Bottom right: High Koenigsberger’s values (Q) obtained from samples from both
Vredefort and Johannesburg domes, South Africa
Johanna Salminen
HELSINKI 2009
Supervisor:
prof. Lauri J. Pesonen
Department of Physics
Division of Atmospheric Sciences and Geophysics
University of Helsinki
Helsinki, Finland
Pre-examiners:
prof. David Evans
Department of Geology and
Geophysics
Yale University
New Haven, United States
Dr. Sergei Pisarevsky
School of GeoSciences
The University of Edinburgh
Edinburgh, United Kingdom.
Opponent:
Custos:
prof. Sten-Åke Elming
Department of Environmental
Engineering
Division of Applied Geophysics
Luleå University of Technology
Luleå, Sweden
prof. Lauri J. Pesonen
Department of Physics
Division of Atmospheric
Sciences and Geophysics
University of Helsinki
Helsinki, Finland
Report Series in Geophysics No. 59
ISBN 978-952-10-5433-4 (paperback)
ISSN 0355-8630
Helsinki 2009
Yliopistopaino
ISBN 978-952-10-5434-1 (pdf)
http://ethesis.helsinki.fi
Helsinki 2009
Helsingin yliopiston verkkojulkaisut
2
PALEOMAGNETIC AND ROCK MAGNETIC STUDY WITH
EMPHASIS ON THE PRECAMBRIAN INTRUSIONS AND IMPACT
STRUCTURES IN FENNOSCANDIA AND SOUTH AFRICA
Johanna Salminen
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 CK112 of Exactum, Gustaf Hällströmin
katu 2b, on May 8th, 2009, at 12 o’clock noon.
Helsinki 2009
3
4
Table of Contents
Abstract .......................................................................................................................8
Acknowledgements ......................................................................................................11
1. Introduction .............................................................................................................12
1.1 Historical review of continental drift ....................................................................12
1.2 Earth’s magnetic field and paleomagnetism as a tool to study continental drift and
supercontinents..........................................................................................................13
1.3 Why dykes and impact rocks?..............................................................................17
1.4 Supercontinents - Rodinia and Hudsonland ..........................................................17
1.4.1 Rodinia .........................................................................................................18
1.4.2 Hudsonland...................................................................................................21
1.5 Goals of this study ...............................................................................................22
2. Methods....................................................................................................................23
2.1 Sampling .............................................................................................................23
2.2 Basic petrophysical measurements .......................................................................24
2.3 Paleomagnetic measurements – used demagnetization techniques........................25
2.3.1 Alternating field demagnetization..................................................................25
2.3.2 Thermal demagnetization ..............................................................................26
2.3.3 Analysis of remanent magnetization components ..........................................26
2.4 Identification of remanence carriers .....................................................................27
2.4.1 Thermomagnetic measurements ....................................................................27
2.4.2 Magnetic hysteresis properties ......................................................................29
2.4.3 Lowrie-Fuller- , Lowrie- and viscosity test....................................................31
2.4.4 Anisotropy of susceptibility ..........................................................................33
2.4.5 Optical methods – scanning electron microscopy studies...............................34
2.5 Paleomagnetic direction and field tests.................................................................34
2.6 Paleomagnetic pole ..............................................................................................36
2.7 Representing continental drift and building up the supercontinents ......................38
3. Summary of the articles I - IV.................................................................................41
3.1 Article I ...............................................................................................................41
3.2 Article II ..............................................................................................................43
3.3 Article III............................................................................................................. 44
3.4. Article IV ...........................................................................................................46
3.5 Author’s contribution...........................................................................................48
4. Updated discussion and synthesis of the results .....................................................49
4.1 New paleomagnetic results for Baltica .................................................................49
4.1.1 Paleomagnetic results from the Valaam sill and from the Salla diabase dyke –
the proposed long-lasting connection between Baltica and Laurentia during the
Mesoproterozoic ....................................................................................................50
4.1.2 Paleomagnetic results of the Jänisjärvi impact structure ................................54
4.2 Vredefort impact structure ...................................................................................57
4.2.1 Paleomagnetic results of the Vredefort impact structure – implication for the
paleoreconstruction of the Kaapvaal Craton at 2.0 Ga................................................57
4.2.2 Younger ca. 1.1 Ga magnetization.................................................................59
4.2.3 Archean magnetization..................................................................................60
4.2.4 High Q values and the origin of remanent magnetization...............................61
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5. Conclusions ..............................................................................................................61
6. References ................................................................................................................63
Original articles ...........................................................................................................81
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This thesis is based on the following four articles, which are refered to in the text by their
Roman numerals:
I
Salminen, J., Donadini, F., Pesonen, L.J., Masaitis, V.K., and Naumov, M., 2006.
Paleomagnetism and petrophysics of the Jänisjärvi impact structure, Russian
Karelia. Meteoritics and Planetary Science, 41, 1853-1870.
II
Salminen, J., and Pesonen, L.J., 2007. Paleomagnetic and rock magnetic study of
the Mesoproterozoic sill, Valaam island, Russian Karelia. Precambrian Research,
159, 212-230.
III
Salminen, J., Pesonen, L.J., Mertanen, S., Vuollo, J., and Airo, M.-L., 2009.
Palaeomagnetism of the Salla Diabase Dyke, northeastern Finland and its
implication to the Baltica – Laurentia entity during the Mesoproterozoic.
Manuscript to be published in: Reddy, S.M., Mazumder, R., and Evans, D.A.D.,
(eds.) Palaeoproterozoic Supercontinents and Global Evolution. Geological
Society of London Special Publication 323.
IV
Salminen, J., Pesonen, L.J., Reimold, W.U., Donadini, F., and Gibson, R.L., 2009.
Paleomagnetic and rock magnetic study of the Vredefort impact structure and the
Johannesburg Dome, Kaapvaal Craton, South Africa – implications for the
apparent polar wander path of the Kaapvaal craton during the Mesoproterozoic.
Precambrian Research, 168, 167-184.
Article I is reprinted from Meteoritics and Planetary Science with permission from the
the Meteoritical Society. Articles II and IV are reprinted from Precambrian Research
with the permission from Elsevier. Article III is reprinted from Geological Society of
London Special Publication with permission from the Geological Society.
7
Abstract
The importance of supercontinents in our understanding of the geological evolution of the
planet Earth has been recently emphasized. The role of paleomagnetism in reconstructing
lithospheric blocks in their ancient paleopositions is vital. Paleomagnetism provides two
important aspects related to supercontinent research. First, it is the only quantitative tool
for providing ancient latitudes and azimuthal orientations of continents. Second, it yields
information of content of the geomagnetic field in the past. In order to obtain a
continuous record on the positions of continents, dated intrusive rocks like diabases are
required in temporal progression. This is not always possible due to pulse-like
occurrences of dykes. In this work we demonstrate that studies of meteorite impactrelated rocks may fill some gaps in the paleomagnetic record.
This dissertation is based on paleomagnetic and rock magnetic data obtained from
samples of the Jänisjärvi impact structure (0.68 Ga, Russian Karelia), the Salla diabase
dyke (1.12 Ga, North Finland), the Valaam monzodioritic sill (1.46 Ga, Russian Karelia),
and the Vredefort impact structure (2.02 Ga, South Africa). The advantage of studying
samples from impact structures is that they may fill the gaps of paleomagnetic data for
building up the apparent polar wander paths (APWP) for continents. In particular, when
the magnetically stable impact rocks can be accurately dated and their natural remanent
magnetization (NRM) passes the impact test, they may provide new paleomagnetic poles
for continents.
The paleomagnetic study of Jänisjärvi impact samples (most recent 40Ar-39Ar age of
682±4 Ma; Jourdan et al., 2008) was made in order to obtain a pole for Baltica, which
lacks paleomagnetic data from 750 to ~600 Ma. The position of Baltica at ~700 Ma is
relevant in order to verify whether the supercontinent Rodinia was already fragmented.
Based on thermal and alternating field demagnetization the impact melt samples yield a
characteristic remanent magnetization (ChRM) of D = 102° and I = 73° (95 = 6.2°). The
ChRM passes the impact test as unshocked basement rocks yield a typical Svecofennian
direction whereas shocked basement rocks revealed similar directions of magnetization as
the impact rocks. The nature of the ChRM is most likely thermal remanent magnetization
(or thermochemical remanent magnetization). Such a direction provides a virtual
geomagnetic pole position of Plat = 45.0°N, Plon = 76.9°E (A95 = 10.4). This pole differs
from the Neoproterozoic poles of Baltica such as the ~750 Ma mean pole and the ~616
Ma pole.
We studied five possible reasons for deviation of the Jänisjärvi pole from Neoproterozoic
poles of Baltica: (1) post-impact remagnetization; (2) post-impact tilting of the area; (3)
secular variation of the geomagnetic field; (4) a poorly defined APWP of Baltica due to
the paucity of well-dated Neoproterozoic paleomagnetic poles or (5) problems with
isotopic dating of impact melt samples. Based on recent 40Ar-39Ar dating (Jourdan et al.,
2008) and our new pole we prefer explanation (4) and a revised APWP for Baltica will be
proposed.
8
The paleomagnetic study of the Salla diabase dyke (U-Pb 1122±7 Ma) located in northeastern Finland was conducted to examine the position of Baltica at the onset of
supercontinent Rodinia's formation.
Paleomagnetic and rock magnetic results are presented from 13 sites sampled along the
dyke. A positive baked-contact test proves that the dyke carries a primary NRM with the
following direction: D = 42°, I = 73° (95 = 4.8°) corresponding to a virtual geomagnetic
pole (VGP) of Plat = 71°N, Plon = 113°E (A95 = 8.1°). Although secular variation may
not have been fully averaged out, the new VGP is far away from any other known
Proterozoic paleopoles of Baltica, thus requiring a revision for the APWP of Baltica. The
pre-Sveconorwegian (~ 1.3 -1.0 Ga) APW swathes of Baltica, Laurentia (including the
Logan Loop) and Kalahari cratons show similar loop-like shapes suggesting that they
may have drifted together. True polar wander is another possible explanation for the
similar loop-like APW shapes (Evans, 2003). However, dated paleomagnetic poles for ca.
1.25 - 1.12 Ga interval from these continents are required to test the similarity.
Nevertheless, the pole from Salla provides hints that the Mesoproterozoic Baltica –
Laurentia unity in the Hudsonland (Columbia, Nuna) supercontinent assembly may have
lasted until 1.12 Ga. In this pre-Rodinia configuration of Baltica and Laurentia, Baltica is
90° rotated compared to Rodinia models of Hoffman (1991), Pesonen et al. (2003) and Li
et al. (2008), but is consistent with configuration of Evans (2006a). The new VGP of
Salla dyke provides new constraint on the timing of the rotation of Baltica relative to
Laurentia (e.g. Gower et al., 1990).
A paleomagnetic study of the well dated (U-Pb 1458+4/-3Ma) Valaam monzodioritic sill
located in NE part of the Fennoscandian Shield was carried out in order to shed light into
the question of existence of Baltica-Laurentia unity in the proposed Mesoproterozoic
supercontinent Hudsonland. As the the Valaam sill is a single intrusive body with a
thickness of 130-150 m, it has been cooled down below 580°C in much less time (order
of hundreds of years as calculated after Jaeger, 1957; and Delaney, 1987) than required to
average out secular variation (order of 104–105 years). Thus the the obtained ChRM of D
= 43°, I = -14° (95 = 3.3°) yields a VGP of Plat = 13.8°, Plon = 166.4°, and A95 = 2.4°.
Combined with results from dyke complex of the Lake Ladoga region (Schehrbakova et
al., 2008) a new robust paleomagnetic pole for Baltica is obtained (Plat = 12°, Plon =
173°, A95 = 7.5°), which places Baltica on a latitude of 10°. This low latitude location is
supported also by Mesoproterozoic 1.5–1.3 Ga red-bed sedimentation (for example the
Satakunta sandstone). According to this study the existence of a single Baltica-Laurentia
block since 1.76 Ga until 1.27 Ga is paleomagnetically permissible and suggests the
possibility that the Hudsonland configuration lasted throughout the long interval.
The ChRM (D = 18°, I = 55°, 95 = 8.1°) of the Vredefort impactite samples passes the
impact test and provides a well dated (2.02 Ga) VGP of Plat = 25.4° N and Plon = 43.8°
E (A95 = 9.6°) for the APWP of the Kaapvaal Craton. Rock magnetic and petrophysical
data reveal unusually high Koenigsberger ratios (Q values) in all pre-impact lithologies,
in some Vredefort impactite samples, and in the much younger (1.1 Ga) gabbro samples.
The high Q values, which have been reported by previous studies of Vredefort
lithologies, are now also seen in samples from the Johannesburg Dome (ca. 120 km
9
away) where there is no impact evidence. Thus, a direct causative link of high Q values to
the Vredefort impact event can be ruled out. As the observed magnetization has high
coercivity and shows multiple components of remanent magnetization, we exclude
lightning as a cause for the high Q values (except in case of gabbros). Instead, the cause
of the high Q values may be related to the circulation of hot fluids taking place within the
two domes (Vredefort and Johannesburg). In the case of Vredefort, the impact event is
almost certainly the triggering cause for the uplifting (doming) of the middle crustal rocks
with enhanced (high Q) magnetization. In the case of Johannesburg, another explanation
for the doming must be sought.
10
Acknowledgements
I wish to thank several people without whom this thesis would have never been
completed:
My supervisor prof. Lauri Pesonen who gave me an opportunity to work in the Solid
Earth Geophysics Laboratory. I appreciate the independence, responsibilities and good
advises he has given me during the project.
Prof. David Evans and Dr. Sergei Pisarevsky provided constructive and critical comments
while reviewing the manuscript of the thesis. Their work is highly appreciated.
Prof. Uwe Reimold is warmly thanked for his co-operation during the Vredefort project.
He introduced me to the beautiful Vredefort impact structure (South Africa). Moreover,
I’ll never forget the hospitality of his and his family during my stay in Johannesburg.
Discussions with Dr. Satu Mertanen, prof. Martti Lehtinen, Dr. Jüri Plado and Dr. Fabio
Donadini enlightened several paleomagnetic and geological problems during this journey.
Their patient help was highly valuable.
Dr. Viktor Masaitis Prof. Roger Gibson, Dr. Alex Deutsch, Mr. Matti Leino, Dr. Jouni
Vuollo, Dr. Meri-Liisa Airo, Dr. Marja Lehtonen, and Ms. Mary Sohn are thanked for
their helpful cooperation during this PhD project.
My colleagues at the Division of Atmospheric Sciences and Geophysics, especially Selen
Raiskila, Tomas Kohout and Tiiu Elbra, are thanked. It has been inspiring to work with
such good people.
This work was funded by the University of Helsinki, Magnus Ehrnrooth’s foundation,
Jenny and Antti Wihuri Foundation, The Academy of Finland, Kordelin foundation and,
Sohlberg’s delegation.
Finally, my friends and family have supported me through this long process. Special
thanks to Anna with whom I was able to share the joy and distress of preparing a PhD
thesis. The help of my mother was invaluable, especially for taking care of my daughter
while I was working. My warmest thanks are addressed to my husband, Veikka, who
believed in me and stood next to me during this whole time.
11
1. Introduction
1.1 Historical review of continental drift
Already in the 17th century, philosopher Sir Francis Bacon (1561-1626) recognized that
the coastline of Africa and South-America matched. Snider-Pellegrini (1802-1885) and
Wegener (1880-1930) further developed this idea and based on known paleontological
evidence (Paleozoic and Mesozoic fossils in South America, Africa, India and Australia
were all very similar), they came up with the proposition that these continents had been
together but later separated (Fig. 1).
Figure 1. In 1858, geographer Antonio Snider-Pellegrini composed these two maps to
explain how the American and African continents may have once fit together,
then later separated. Left: The formerly joined continents before separation.
Right: The continents after separation. b) As noted by Snider-Pellegrini and
Wegener, the locations of certain fossil plants and animals on present-day,
widely separated continents would form definite patterns (shown by the bands),
if the continents were rejoined.
Before Wegener, the connections were considered to be former land bridges which
subsequently sank to form the ocean basins. Based on the gravity data over the oceans,
Wegener proposed in 1912 that the ocean floor was not just a sunken continent; it was
made of entirely different composition (e.g. Jacoby, 2001). Wegener introduced the
concept of horizontal ”continental drift,” suggesting that the continents move around the
Earth independently. While moving, continents close oceans in front of them and open
new oceans behind them. The mobile asthenosphere was found to be overlain by a rigid
lithosphere that contained both upper mantle and crust, indicating that continental drift
implied drift of the lithosphere (e.g. Vine, 1966; Morgan, 1968; LePichon, 1968). Later
paleomagnetic information demonstrated that different continents move separately (e.g.
Creer et al., 1954). After seafloor spreading (Dietz, 1961; Hess, 1962) was recognized
and spherical trigonometrical rules of plate tectonics were developed (Le Pichon, 1968;
Morgan, 1968), the concept of plate tectonics was concieved.
12
According to Cawood et al. (2006), several lines of evidence demonstrate that plate
tectonics has been an active component of the Earth’s processes since the formation of
the first continental crust at >4.3 Ga: (1) independent lateral motion of lithospheric
blocks, which is supported by paleomagnetic data; (2) magmatic arc activity and
associated ore deposits related to subduction of oceanic-type lithosphere, which is
supported by geochemical data; (3) fossil subduction zones obtained by seismic imaging;
and (4) assembly of continental lithosphere along linear orogenic belts indicated by
tectonostratigraphic associations. However, by evaluating rocks created in subduction
zones (e.g. ophiolites, blueschists, and ultrahigh pressure metamorphic terranes) Stern
(2005) proposed that subduction tectonics began only in the Neoproterozoic era.
Recently, by demonstrating differential lateral motion between internally rigid
lithospheric blocks older than 1110 Ma, Evans and Pisarevsky (2008) quantitatively
refuted the hypothesis of Stern (2005). Using three different paleomagnetic approaches,
Evans and Pisarevsky (2008) identified differential motion of cratons as old as 2445–
2680 Ma. According to this evidence, modern-style plate tectonics appears to have been
in operation throughout most of the Proterozoic. They also have evidence of activity as
early as the Archean.
Supercontinents are assemblies that contain a large proportion of the Earth’s continental
blocks (e.g. Rogers and Santosh, 2004). Evidence supports that supercontinents have
played a key role in the evolution of the Earth’s surface, possibly since Archean times,
but as Evans and Pisarevsky (2008) demonstrated at least since Proterozoic times. The
supercontinent cycle – amalgamation and subsequent break up of supercontinents have
had major impact on geology and biology, such as forming the orogenic belts,
diversification of life forms on Earth, unique climatic conditions (glaciations), global
changes in ocean chemistry, long-lived mantle convection patterns giving rise to plumes
and large igneous provinces, and variations in heat-flow across the core-mantle-boundary
(e.g. Condie, 2000; 2004; Karlstrom et al., 2001; Meert, 2002; Rogers and Santosh, 2002;
Pesonen et al., 2003a; Zhao et al., 2003; 2004 and references therein; Evans and
Pisarevsky, 2008).
1.2 Earth’s magnetic field and paleomagnetism as a tool to study continental drift
and supercontinents
In this study, the paleomagnetic method, the only quantitative tool to measure lateral
motion of continental blocks, is used to study the ancient continental drift. Rocks acquire
their magnetization in Earth’s magnetic field (domains of magnetic minerals align during
the cooling of a rock consistently with the direction of the geomagnetic field at that
particular time), which can be measured in a laboratory. Earth’s magnetic field is
believed to be driven by dynamo processes and that the polarity of this magnetic field
changes a few times in a million years. However, several superchrons have been observed
– long periods when no reversals take place (for example the Cretaceous Long Normal,
the Jurassic Quiet Zone and the Kiaman Long Reversed). When averaged over a
13
sufficient time interval (~ 105 - 106 years), the Earth’s magnetic field may be
approximated by geocentric axial dipole (GAD) (Hospers, 1954; Irving, 2005).
Origin of the Earth’s magnetic field is thought to be some form of a self-sustaining
dynamo, where the solid inner core is rotating inside the liquid outer core. Fluid motions
in the Earth’s core are sufficient to regenerate the magnetic field, but there is enough
leakage to keep the shape of the geomagnetic field fairly simple. Thus, the dominant
portion of the geomagnetic field is dipolar, especially at the Earth's surface where, 3000
km above the source, smaller-scale features are filtered away. Fluid eddy currents within
the core –near the core-mantle boundary – produce some subsidiary nondipolar features
of the geomagnetic field. However, nondipole components of geomagnetic field may
cause bias in paleomagnetic data (Pesonen and Nevanlinna, 1981; Meert et al, 2003). For
example, the magnitude of the octupole component has been proposed to vary between
negligible levels and as much as ~20% relative to the dipole during the past 300 Ma (e.g.
Torsvik and Van der Voo, 2002). The majority (99%) of Earth’s magnetic field is caused
by this internal origin. Some external sources also contribute to the Earth’s magnetic
field, such as the Sun’s activity and associated ionospheric and magnetospheric electric
currents, but these sources are not discussed in this thesis. In some cases, lightning may
magnetize crustal rocks. Some crustal magnetic sources also contribute to the Earth’s
magnetic field and they may affect the paleomagnetic measurements if sampled (for
example causing errors when orienting samples).
The magnetic dip pole (defined as the point on the Earth’s surface where the field is
perfectly vertical; I = ±90°) wanders with a speed of a few degrees per century around the
geographic North Pole, caused by the secular variation of the geomagnetic field (Merrill
et al., 1996). Over geological times (105 to 106 years) the secular variation is averaged out
and the magnetic field can be represented by a geocentric axial dipole model (GAD)
shown in Fig 2a. In this model, the magnetic field is produced by a single magnetic
dipole at the centre of the Earth and aligned with the rotation axis (e.g. Butler, 1992).
Using a GAD the ancient latitude (paleolatitude) of the sampling site may be obtained
from the magnetic inclination, applying the basic dipole equation:
tanI = 2tan
(1)
where I stands for inclination and for paleolatitude. I increases from –90° at the
geographic south pole to +90° at the geographic north pole. Lines of equal I are parallel
to lines of latitude and are simply related through Equation (1). For a GAD, D = 0°
everywhere.
Inclinations of the present geomagnetic field are positive in the northern hemisphere and
negative in the southern hemisphere and the geomagnetic equator (line of I = 0°) is close
to the geographic equator. The magnetic dip poles are not at the geographic poles as
expected for a GAD field, and the magnetic equator wavers about the geographic equator.
In a modification of a GAD model, a geocentric dipole is inclined to the rotation axis, as
14
shown in Fig. 2b. The inclined geocentric dipole that best describes the present
geomagnetic field has an angle of ~11.5° with the rotation axis. The poles of the bestfitting inclined geocentric dipole are the geomagnetic poles, which are points on the
surface where extensions of the inclined dipole intersect the Earth’s surface. If the
geomagnetic field were exactly that of an inclined geocentric dipole, then the
geomagnetic poles would exactly coincide with the dip poles. The fact that these poles do
not coincide indicates that the geomagnetic field is more complicated than can be
explained by a dipole at the Earth’s centre (e.g. Butler, 1992).
a
b
Figure 2. a) Model of a geocentric axial dipole. In this model magnetic dipole M is
placed at the centre of the Earth and aligned with the rotation axis; the
geographic latitude is ; the mean Earth radius is re; the magnetic field
directions at the Earth’s surface produced by the geocentric axial dipole are
schematically shown; inclination, I, is shown for one location; N is the north
geographic pole. b) Inclined geocentric dipole model. The best-fitting inclined
geocentric dipole is shown in meridional cross section through the Earth in the
plane of the geocentric dipole; distinctions between magnetic poles and
geomagnetic poles are illustrated; a schematic comparison of geomagnetic
equator and magnetic equator is also shown. Redrawn by Butler (1992) after
McElhinny (1973).
The GAD-model has been the basic model for paleomagnetic application since the 1950’s
(see Irving, 1964). However, its existence during the Precambrian times must be tested
(e.g. Evans, 2006b; Korhonen et al., 2008). The possible tests are as follows: (1) Study
the origin of observed magnetization. (2) Study the paleoclimate latitude indicators such
as coral reefs (shallow latitudes), evaporites (shallow to moderate latitudes) and
glaciogenic rocks (high latitudes) (Evans, 2006b). (3) Testing directly the dipole equation
(tanI = 2tan) by comparing the observed inclination and intensity curves as a function of
15
latitude with the ones provided by the GAD. (4) Plotting the frequency distribution of
inclinations (Evans, 1976). (5) Other tests such as testing the reversal asymmetries
(Pesonen and Nevanlinna, 1981), or the comparison of the latitudinally dependent
paleosecular variation curves with some prescribed field model (e.g. Model G, Merrill
and McElhinny, 1983) for both polarities and hemispheres (Smirnov and Tarduno, 2004).
Remanent magnetization as old as ca. 3.5 Ga is observed from the Komati Formation in
South Africa (Hale and Dunlop, 1984; Yoshihara and Hamano, 2004) and from the
Duffer Formation in North West Australia (supported by positive fold test; McElhinny
and Senanayake, 1980). This reinforces that the behavior of the magnetic field and
dynamo has been similar to present day since Archean times. Moreover, studies of
Paleoproterozoic intrusions (e.g. ca. 2.5 Ga Burakovka intrusion, Russia; Smirnov et al.,
2003) suggest that the intensity of the geomagnetic field was similar to the present-day
intensity. In contrast, Shcherbakova et al. (2008) questioned the validity of a GAD for the
Archean, Proterozoic and starting time of dynamo by presenting and summarizing studies
showing low paleointensity values in the Archean and Early Proterozoic subsequently
followed by higher values in Late-/Middle- Proterozoic.
Support for dynamo activity since the Late Archean-Early Proterozoic began when
Smirnov and Tarduno (2004) pointed out that the magnitude of secular variation at ca. 2.5
Ga (ca. 2.5 Ga dykes from Superior and Karelia) was already similar to the present one.
Recent study by Biggin et al. (2008) shows that the dynamo was already working at 2.5
Ga, but contradictory to Smirnov and Tarduno (2004), on average, geomagnetic secular
variation during the late Archaean and early Proterozoic was different from that of the
past 200 million years. Studies of the rocks from the Kaapvaal craton (the Dalmein and
Kaap Valley Plutons that intrude the Barberton greenstone belt) indicate the start of the
dynamo between 3.9 and 3.2 Ga (Tarduno et al., 2007; see also Dunlop, 2007). Moreover,
Evans (2006b) showed that the GAD model fits the distribution of evaporite basins on
Earth throughout most of the last two billion years (relevant for articles I, II and III). For
> 2 Ga, the evaporites are less voluminous, but there is an excellent consistency among
many rock recorders of the apparent polar wander path of Kaapvaal inbetween 2.06-1.93
Ga (e.g. de Kock et al., 2006 and references therein; relevant for article IV) pointing to
long-term stability at least for moderate-sized areas of the Earth's surface over 100million-year intervals.
On longer time scales, the dipolar geomagnetic field has been observed to switch polarity.
The present configuration of the geocentric dipole (pointing toward geographic south) is
referred to as normal polarity. The opposite configuration is defined as reversed polarity.
Reversal of the dipole’s polarity produces (in most cases) a 180° change in surface
geomagnetic field direction at all points. However, departures of 180° symmetry are
occasionally documented (e.g. Pesonen and Nevanlinna, 1981 and references therein).
Reversals were first noted, from the spreading sea floor, as magnetic anomalies over
oceanic ridges (Cox et al., 1963; Vine and Matthews, 1963).
Paleomagnetic studies of Paleo-Mesoproterozoic rocks must distinguish primary from
secondary magnetizations in order to be useful. In addition, the age of the paleomagnetic
16
poles must be known with reasonable certainty if they are to be used for reliable
paleoreconstructions (e.g. Van der Voo, 1990; Van der Voo and Meert, 1991; Buchan et
al., 2000). Van der Voo (1990) formulated the following seven criteria for reliable
paleomagnetic data: (1) Data is obtained from samples with well-determined rock age
and a presumption that the magnetization age is the same age. (2) The number of samples
for the study (n>24) is sufficient. (3) Adequate demagnetization technique that
demonstrably includes vector subtraction, or principal component analysis, is used. (4)
The age of magnetization should be constrained by field tests. (5) There should be
structural control and tectonic coherence with the block or craton involved. (6) Evidence
of the presence of reversals is documented. (7) There should not be resemblance to
paleomagnetic poles of a younger age (by more than a period).
There are several tools for reconstructing supercontinents (e.g. Rogers and Santosh,
2004). However, the only quantitative method for testing supercontinents is
paleomagnetism (e.g. Meert, 2002; Evans and Pisarevsky, 2008). The method is
described in section 2.7 of this thesis. The other methods are based on geology and
include: (1) correlation of orogenic belts that developed during accretion of the
supercontinent; (2) correlation of extensional features that developed when the
supercontinent fragmented; and (3) recognition that the source of sediment in continent A
is within continent B supports the idea that continent B was earlier adjacent to continent
A (e.g. Meert, 2002).
1.3 Why dykes and impact rocks?
Dykes and dyke swarms hold some of the keys to the interpretation of plate tectonics as
they provide information on the extensional processes occurring both in the continental
and oceanic lithosphere. Mafic dyke episodes are important geological time markers that
provide a tool for continental reconstructions as they are amenable for precise dating and
preserving a stable record of ancient magnetic fields (e.g. Buchan and Halls, 1990).
Moreover, dykes are easy to recognize and sample. To obtain a continuous record on
positions of continents, dated intrusive rocks are required in temporal progression, which
is not always possible due to pulse-like occurrences of dykes. In this work, we
demonstrate that studies of impact-related rocks may fill some gaps in the paleomagnetic
record. In particular, when the magnetically stable impact rocks can be accurately dated
and if their natural remanent magnetization (NRM) passes the impact test (Pesonen,
2001), they may provide additional paleomagnetic poles for continents. Impactites may
also provide material for paleointensity studies (Nakamura and Iyeda, 2005; see also
Article I).
1.4 Supercontinents - Rodinia and Hudsonland
Supercontinents have played a key role in the evolution of the Earth’s surface most
probably since Archean times. There appear to have been several times in Earth history
when the continents were fused into large masses or when accretion formed smaller
17
blocks that contained parts of more than one continent (e.g. Dalziel, 1995; Rogers, 1996;
Rogers and Santosh, 2003, Pesonen et al., 2003a, Li et al., 2008). The concept of three
youngest supercontinents Pangea (Pangaia), Gondwana and Rodinia are widely accepted
(e.g. Bond et al., 1984; McMenamin and McMenamin, 1990; Dalziel, 1995; Evans,
2006a). There is ample evidence for existence of the youngest supercontinent – Pangea.
Reconstruction of Pangea suggests that it assembled through plate convergence at the
expense of now vanished Neoproterozoic–Paleozoic oceans and that it contained almost
all of the world’s continental crust (e.g. Evans, 2006a). Gondwana assembled in
Neoproterozoic-Cambrian time via numerous collisions.
Figure 3. Three supercontinents Pangea (Pangaia), Rodinia and Hudsonland (after
Pesonen and Sohn, in press).
1.4.1 Rodinia
Nowadays it is widely accepted that Rodinia assembled through worldwide orogenic
events between 1300 and 900 Ma. The variations of reconstructions of Rodinia show a
network of Archean cratonic nuclei, Paleoproterozoic foldbelts, and early
Mesoproterozoic rift-to-passive-margin cover successions, suggesting also that Rodinia
assembled via amalgamation of fragments from an earlier supercontinent (Evans, 2006a).
In most studies Laurentia is thought to have formed the core of Rodinia since it is
surrounded by Neoproterozoic passive margins (e.g. Bond et al. 1984; McMenamin and
McMenamin, 1990; Hoffman, 1991; Meert and Torsvik, 2003; Li et al., 2008).
Otherwise, there are several competing models of configuration and the life-time of
Rodinia. The configuration proposed by Hoffman (1991) assumes that all Grenville-age
(ca. 1300-1000 Ma) orogenic belts are zones of oceanic closure between continental
blocks (Fig. 4a). It allows a juxtaposition of western Laurentia, Australia and East
Antarctica referred to as SWEAT (Southwest U.S. - East Antarctic) proposed by Moores
(1991) with the corollaries by Dalziel (1991) and by Hoffman (1991). Powell et al.
(1993) confirmed that the SWEAT connection was paleomagnetically permissible for the
time interval of ca. 1050-720 Ma based on the contemporary paleomagnetic database.
However, later Wingate and Giddings (2000) refuted it at 755 Ma, and Wingate et al.
(2002) at 1070 Ma. Furthermore, Pisarevsky et al. (2003a) proposed that such a
18
connection is not possible at ca. 1200 Ma. Besides, Li et al. (2008) list several preGrenvillian geological mismatches concerning the SWEAT connection.
Figure 4. Different Rodinia models: a) SWEAT configuration by Hoffman (1991); b)
AUSMEX configuration of Laurentia and Australia in Rodinia by Wingate et
al. (2002); c) Configuration of Li et al. (2008) based on “missing-link”-model
(Li et al., 1995); and d) Rodinia by Evans (2006a).
The matching of basement provinces and sedimentary provenances between Australia
and south-western Laurentia made by Karlstrom et al. (1999; 2001) and by Burrett and
Berry (2000, 2002) generates the AUSWUS connection originally proposed by
Brookfield (1993). Li et al. (2008) indicated several geological mismatches also
concerning this model, such as the lack of a prominent ca. 1400 Ma granite-rhyolite
province in Australia as present in southern Laurentia, the scattered evidence of
Grenvillian metamorphism, mismatches in ca. 825 Ma and 780 Ma plume records, and
starting age of continental rifting between these continents. Moreover, Wingate et al.
(2002) negated AUSWUS model at 1070 Ma and later Pisarevsky et al. (2003a) showed
that there also is ca. 1200 Ma paleomagnetic misfit concerning this configuration.
Li et al. (1995) proposed the “Missing-Link” model, where the South China Block is
reconstructed between Australia-East Antarctica and Laurentia due to the mismatches in
19
the crustal provinces of Australia-East Antarctica and Laurentia in the SWEAT
connection, similarities in the Neoproterozoic stratigraphy of South China, southeastern
Australia and western Laurentia, and the need for western source region to provide the
late Mesoproterozoic detrial grains in the Belt Basin. Li et al. (2008) adapted this
configuration for constructing a Rodinia map (Fig. 4c). This model negates the need for
matching basement geology between Australia-East Antarctica and western Laurentia
prior to 1000-900 Ma. It is also supported by the Neoproterozoic rift record, the early
Neoproterozoic mantle plume record, and paleomagnetic constraints (Li et al., 2008).
However, this model is not long-lived as it assembled as late as 900 Ma and began to
fragment as early as 825 Ma.
Based on well defined paleomagnetic poles from the Bangemall Basin sills in the
Edmund Fold Belt of Western Australia Wingate et al. (2002) proposed the AUSMEX
(northern Queensland of Australia against Mexico of southern Laurentia) model (Fig. 4b).
This model also negates the need for matching basement geology between Australia and
Laurentia, and it aligns the Grenville belt with coeval orogens in Australia. Moreover,
paleomagnetic pole from the 800 Ma to 760 Ma oriented cores from the Lancer-1 drill
hole in the Officer Basin of south Central Australia is more compatible with the Ң780 Ma
Laurentian poles in the AUSMEX configuration than in any of the other (e.g. AUSWUS,
SWEAT, Missing-Link) configurations (Pisarevsky et al., 2007). However, Pisarevksy et
al. (2003a) negate the AUSMEX juxtaposition at 1200 Ma.
By analyzing paleomagnetic data for different Rodinia reconstructions Evans (2006a),
postulated that recent paleomagnetic data could require that Rodinia was rather small or
short-lived than the great and enduring early Neoproterozoic supercontinent as originally
conceived. Based on reliable paleomagnetic poles and global tectonostratigraphy Evans
(2006a) presented a new radically revised Rodinia model allowable for inclusion of all
the largest 10-12 cratons spanning the entire time interval of ca.1100–750 Ma, i.e., in
accordance with the original concept of a supercontinent assembled via "Grenvillian"
collisions and disaggregated by mid-Neoproterozoic rifts (Fig. 4d). The new Rodinia
model of Evans (2006a) includes few traditional cratonic connections – the only
successful paleomagnetic confirmation of a tectonically-based reconstruction is that of
northeast Laurentia and Baltica (Pesonen et al., 2003a; Pisarevsky et al., 2003b).
The relative position of Baltica to Laurentia in Rodinia is one of the best known.
Geological evidence shows that the two blocks could have been together as early as ca.
1800 Ma (e.g. Gower et al., 1990; Gorbatschev and Bogdanova, 1993; Karlstrom et al.,
2001; Evans, 2006a). Earlier paleomagnetic data suggest a rather complex history
between the two cratons during the Mesoproterozoic (e.g. Elming et al., 1993, 2001;
Buchan et al., 2000, 2001; Pesonen et al., 2003). The configuration adopted in the
Rodinia model of Li et al. (2008) in Fig. 4 follows that of Hoffman’s (1991) geologically
based reconstruction. This “right-side-up” Baltica connection with Laurentia is the least
controversial – agreed upon by both Evans (2006a) and Li et al. (2008), but it should be
re-evaluated because of the non-coincident ages of the Sveconorwegian and Grenville
loops. Pisarevsky and Bylund (2006) have shown that the high-paleolatitude
Sveconorwegian data is probably the result of a post-900 Ma regional cooling after
20
metamorphism whereas the apex of the Grenville APW loop for Laurentia is the 1015±15
Ma Haliburton A pole (Warnock et al., 2000). Paleomagnetic data from the rocks older
than ca. 1100 Ma do not permit exactly the same configuration but they still allow the
two adjacent cratons to develop correlative Proterozoic belts including the
Sveconorwegian and Grenville orogens and similar ca. 1500-1350 Ma intracratonic
magmatism (e.g. Li et al., 2008 and references therein).
Several lines of evidence from paleomagnetic, tectonostratigrapic, and geochronology
studies indicate that younger supercontinents assemble via amalgamation of fragments
from an earlier supercontinent (e.g. Meert and Torsvik, 2003, Evans, 2006a). Timing of
the breakup of Rodinia is broadly synchronous with the assembly of Gondwana (e.g.
Meert and Torsvik, 2003; Evans 2006a). According to Li et al. (2008), Rodinia already
started to fragment already at 825 Ma. Due to paleomagnetic results and mechanism for
the formation of the proposed Rodinia superplume, Li et al. (2008) propose that both the
superplume and Rodinia above it may have moved from a high-latitude position to an
equatorial position between ca. 820-800 Ma and ca. 780-750 Ma.
1.4.2 Hudsonland
Evidence for pre-Rodinia supercontinents becomes more controversial as the age of the
geological formations increase, due to paucity of paleomagnetic data and inherent dating
problems. However, most cratonic blocks show evidence of collisional events occurring
between 2.1 and 1.8 Ga, which has led many researchers to propose that Mesoproterozoic
supercontinents, such as Hudsonland (also known as Columbia and Nuna), existed in the
Early Proterozoic (e.g. Meert, 2002; Rogers and Santosh, 2002; Pesonen et al., 2003a;
Zhao et al., 2003; 2004 and references therein; Condie, 2004; Evans, 2006a).
Figure 5. Some proposed Hudsonland (Columbia) supercontinent configurations. a) The
ca.1.5 Ga paleomagnetic reconstruction after Pesonen et al. (2003a). b)
Reconstruction modified from Zhao et al. (2004) based on lithostratigraphic,
tectonothermal, geochronological and paleomagnetic data (only continents,
which have 1.5 Ga paleomagnetic data are shown). c) Reconstruction of
Karlstrom et al. (2001) supporting AUSWUS model.
21
Recently, several reconstructions of assemblies of supercontinent Hudsonland have been
presented. For example, based on paleomagnetic data Pesonen et al. (2003a) have
suggested that Laurentia and Baltica coexisted more than 600 Ma (from ca. 1.83 to ca.
1.20 Ga), forming together with Amazonia, and perhaps with some other continents, such
as Australia, the core of this Early Proterozoic supercontinent (Fig. 5). Several other
authors have suggested a variety of configurations and lifecycles for this supercontinent
(e.g. Karlstrom et al., 2001; Meert, 2002; Zhao, 2004 and references therein; Hou et al.,
2008a, 2008b). The maximum packing of continents at Hudsonland is believed to have
occurred ca. 1.5 Ga ago (e.g. Meert, 2002; Pesonen et al., 2003; Rogers and Santosh,
2002; Zhao et al., 2004 and references therein). According to these authors, the
fragmentation of Hudsonland began ca. 1.5 Ga ago in association with development of
continental rifts and widespread anorogenic activities, which included e.g. emplacement
of anorthosite massifs, charnockite intrusions, rapakivi granites, and other intrusions,
such as the Valaam monzodioritic sill. The debate regarding the time of the final break up
of Hudsonland is ongoing. The time is proposed to be related: (1) to the ca. 1.6–1.4 Ga
bimodal rapakivi magmatism (Rämö and Haapala, 2005); (2) to the rifting and
emplacement of the 1.3–1.2 Ga mafic intrusions (e.g. SW Finland and central Sweden)
and the coeval mafic dyke swarms in North America (e.g. Elming et al., 2001); or (3) to
the intensive pulse of 1.3 – 1.2 Ga large igneous provinces producing giant radiating dyke
swarm (Hou et al., 2008a,b).
1.5 Goals of this study
The main objective of this work was to obtain additional Neoproterozoic and
Mesoproterozoic data for the apparent polar wander path of Baltica and to study the
proposed long-lived (1.82 – 1.27 Ga) connection between Baltica and Laurentia (e.g.
Evans and Pisarevsky, 2008). A paleomagnetic study of the Jänisjärvi impact samples
(most recent 40Ar-39Ar age of 682±4 Ma; Jourdan et al., 2008) was made to obtain a
pole for Baltica, which lacks paleomagnetic data from 750 to ~600 Ma. The position of
Baltica at ~700 Ma is relevant to verify whether the supercontinent Rodinia had already
fragmented. A paleomagnetic study of the Salla diabase dyke (1122 ± 7 Ma; Lauerma,
1995), located in northeastern Finland, was made to study Baltica’s position at the time of
Rodinia's assembly. A paleomagnetic study of the well dated (1458 +4/-3 Ma; Rämö et
al., 2001; Rämö and Haapala, 2005) Valaam monzodioritic sill, located in the
northeastern part of the Fennoscandian Shield, was carried out to shed light on the
questioned existence of the Paleo-Mesoproterozoic supercontinent Hudsonland.
The second objective was to verify the suitability of magnetically stable impactite
samples from two different impact structures (e.g. Jänisjärvi impact structure from
Russian Karelia, and Vredefort impact structure from South Africa) for paleomagnetic
study (Fig. 6).
The third objective was to provide new poles for the Paleoproterozoic apparent polar
wander path of the Kaapvaal craton using samples from the Vredefort impact structure
22
and surroundings. We also tried to determine the original magnetization of the Archean
basement rocks of the Vredefort impact structure.
The fourth objective of this work was to study the unusually high Koenigsberger ratios
(Q values) observed in various lithologies around Vredefort structure. This was done
using numerous rock magnetic experiments.
Figure 6. Paleomagnetic study of dated impact rocks (impact melt) was carried out in
order to verify their suitability for paleomagnetic study. Left: impact melt from
the Jänisjärvi impact structure (Russian Karelia). Photo: V. Masaitis. Right:
Pseudotachylitic (PT) breccia veins from the Vredefort impact structure (South
Africa). Photo: R. Gibson.
2. Methods
2.1 Sampling
Samples for this study have been taken from a ca. 0.7 Ga old (Masaitis et al., 1976;
Müller et al., 1990; Jourdan et al., 2008) Lake Jänisjärvi impact structure (Russian
Karelia), a ca. 1.12 Ga old (Lauerma 1995) Salla Diabase Dyke (northern Finland), a 1.46
Ga old (Rämö et al., 2001; Rämö and Haapala, 2005) Valaam Sill (Lake Ladoga, Russian
Karelia), and from a 2.02 Ga old Vredefort impact structure (South Africa). The majority
of the samples were hand samples and a few were taken with the help of portable field
drills. Samples were oriented by a magnetic or sun compass (Fig. 7). Described in section
3.5 of this work is the author’s contribution to sampling. Multiple specimens were
prepared from the each sample. In case of hand samples we were able to prepare more
specimens (from several drill cores per sample drilled in laboratory) than in case of core
23
samples (only one drill core per sample). Prepared specimens were the size of standard
cylinders (2.5 cm diameter (D), 2.2 cm length (h), h/D = 0.9).
Figure 7. ) Monzodioritic hand sample from Valaam sill. Orienting the samples was done
using the magnetic compass. b) Orienting the drill core of Salla diabase dyke.
The numbers and letters denote marking system. Photos: J. Salminen
2.2 Basic petrophysical measurements
Proper paleomagnetic study requires in-depth knowledge of petrophysical and rock
magnetic properties of the studied samples. Preliminary petrophysical measurements of
natural remanent magnetization (NRM), of magnetic susceptibility () and of density,
were carried out for each specimen before paleomagnetic and other rock magnetic
studies. Petrophysical properties are mainly informative about the remanence stability
(e.g. NRM and Koenigsberger ratio), but they also give some hints about mineralogy and
mineralogical changes (e.g. susceptibility). Curie-point analyses give information about
the magnetic carriers in specimens whereas hysteresis measurements give information
about the domain states of these carriers.
The NRM was measured using a fluxgate magnetometer of the Risto-5 system. The
susceptibility indicates the amount of magnetic material in the specimen and is a measure
of the ease with which the material is magnetized in the presence of an external magnetic
field (magnetizability of a substance). Susceptibility has been measured at room
temperature using the Risto-5 kappabridge, working at frequency of 1025 Hz and a field
48 Am-1. The Koenigsberger ratio (Q value) is defined as the ratio of the remanent
magnetization (a recording of past magnetic fields that have acted on the material) to the
induced magnetization in the Earth’s magnetic field. The International Geomagnetic
Reference Field (IGRF) at a particular site (B) can be used to determine Q value, using
the formula:
24
Q
P 0 NRM
FB
(2)
where Q is Koenigsberger’s ratio, μ0 is the magnetic permeability in vacuum (4×10-7
Hm-1), NRM is natural remanent magnetization (mA/m), is volume susceptibility (10 -6
SI), and B is geomagnetic reference field value (50 μT was used).
According to equation (2) if Q > 1, the remanent magnetization is stronger than the
induced magnetization and the specimen is able to maintain a stable remanence (e.g.
Stacey and Banerjee, 1974). On the other hand, high Q values (> 10) may indicate that
some strong magnetic fields, other than the Earth’s magnetic field (e.g. lightning-induced
field, or shock-produced plasma field see article IV), have affected the rock, and such a
sample is not suitable for paleomagnetic study with tectonic applications.
2.3 Paleomagnetic measurements – used demagnetization techniques
Since in this study we studied samples from dykes, sills and impact structures, we believe
that the primary characteristic remanent magnetization of the samples is of thermal origin
acquired during cooling. However, due to the fact that studied samples are of
Precambrian age such rocks may have acquired several types of secondary magnetization
(e.g. Thermochemical Remanent Magnetization (TCRM), Chemical Remanent
Magnetization (CRM), Viscous Remanent Magnetization (VRM), or Shock Remanent
Magnetization (SRM)). Various components of magnetization sum together to constitute
the vector NRM. The objective of paleomagnetic laboratory work is to isolate various
remanence components and to identify the origin, reliability, and meaning of each
component. Both alternating field (AF) and thermal (TH) demagnetization techniques
were used to separate the characteristic remanence component (ChRM) out of probable
secondary overprints.
The measurement of the remanent magnetization was performed using an AGICO spinner
magnetometer (JR-6a) or a 2G DC-SQUID magnetometer. The spinner magnetometer,
which has a sensitivity of 2 × 10-5 A/m for magnetization and a measuring range up to
12500 A/m, was used for the majority of the specimens from the Lake Jänisjärvi impact
structure (Article I) and strongly magnetic rocks of the Vredefort impact structure
(Article IV). In general, the automatic setting was used, which allows the measurement in
three orthogonal positions. Normally, the high rotation frequency (87.7 Hz) was selected.
Most of the specimens were measured using the SQUID magnetometer, which sensitivity
is 10-6 A/m.
2.3.1 Alternating field demagnetization
25
In the alternating field (AF) demagnetization technique, a specimen is exposed to
progressively increasing alternating magnetic field. The waveform is sinusoidal with a
linear decrease of magnitude with time. AF demagnetization can be used to erase NRM
carried by grains with coercivities less than the used peak demagnetizing field. Used
instruments in this study allowed a maximum field of 160 mT (the 2G-SQUID
magnetometer) or 100 mT (AGICO LDA3 AF-demagnetizer). The AF technique is a
rather fast, cleaning procedure compared to the thermal demagnetization technique.
2.3.2 Thermal demagnetization
When performing stepwise thermal (TH) demagnetization the samples are heated to
elevated temperatures below and around the Curie temperatures of ferromagnetic
minerals and then cooled back to room temperature in zero magnetic field. This causes all
the magnetic grains with blocking temperatures less than the applied temperature to lose
that part of their NRM. After each temperature step, the remaining magnetization and
also the susceptibility was measured. The susceptibility measurement is used to monitor
thermally induced alterations of mineralogy and possible laboratory induced TRMs.
TH measurements were performed using a Schoenstedt TD1 oven (Article I, II, and IV),
the Magnetic Measurement Co thermal demagnetizer (Article II, III, and IV) and the ACS
thermal furnace of the paleomagnetism laboratory of the Otago University in New
Zealand (Article III).
It is necessary to use both AF and TH techniques and compare the results. Some
specimens will be progressively demagnetized using the AF while other specimens using
only the TH technique. The overall objective is to reveal the NRM components which are
carried by ferromagnetic grains within a particular interval of coercivity or blocking
temperature spectra.
2.3.3 Analysis of remanent magnetization components
To isolate different remanence components, standard multicomponent analysing methods
were applied to the data such as vector plots, principal component analysis (Kirschvink,
1980; Leino, 1991) coupled with ordinary vector subtraction technique, and intersecting
great circle-technique with end-point analysis (Halls, 1976; 1978). The majority of the
results were analysed with Leino’s Tubefind program (1991). It allows determination of
the components of directions of magnetization using a least square approach in 3dimensional space and its uncertainty. The software also plots the standard orthogonal
vector plots of Zijderveld (1967), a stereoplot of directions, and a decay of intensity in the
course of demagnetization steps.
26
2.4 Identification of remanence carriers
When evaluating the origin and type of the separated paleomagnetic components it is
important to identify the magnetic carriers of each component. Both AF and TH
demagnetization techniques give hints about the carrier of magnetization (Dunlop and
Özdemir, 1997), but rock magnetic measurements and optical observations of thin
sections are needed to confirm the nature of such minerals. For this work both high and
low temperature thermomagnetic treatments, together with measurements of hysteresis
properties, Lowrie-test (Lowrie, 1990), Lowrie-Fuller test (Lowrie and Fuller, 1971),
studies of anisotropy of susceptibility, and scanning electron microscope analysis were
carried out
2.4.1 Thermomagnetic measurements
In order to understand the acquisition of paleomagnetic recording in rocks, we need to
study the mineralogy of the ferromagnetic minerals within our samples. The fundamental
property of ferromagnetic solids is their ability to record the direction of an applied
magnetic field (e.g. Dunlop and Özdemir, 1997). For a given ferromagnetic material and
temperature there is a maximum magnetization recorded by the sample under applied
field referred to as saturation magnetization (M S). If the solid has acquired its MS, further
increase of the external magnetic field (H) will not result in increased magnetization. At
temperatures below Curie temperature (TC), the magnetic moments are partially aligned
within magnetic domains in ferromagnetic materials. As the temperature is increased
from below the TC, thermal fluctuations increasingly destroy this alignment until the net
magnetization becomes zero at and above the T C (e.g. Dunlop and Özdemir, 1997)
resulting in paramagnetic behavior (paramagnetic solids contain atoms with atomic
magnetic moments, which do not interact between adjacent moments). However, in the
presence of the Earth’s magnetic field, the paramagnetic materials acquire a small
magnetization which is inversely proportional to temperature. At this stage, the
susceptibility of the material follows the Curie-Weiss law and can be expressed as:
F
M
H
C
T
(3)
where M is the magnetization (A/m), H the magnetic field (T), C is the Curie constant,
and T is the temperature (K).
As ferromagnetic minerals have varying TC’s, high temperature thermomagnetic
measurements are useful to investigate the magnetic mineralogy of particular specimens.
In Table 1, the typical TC-values are given for the common minerals encountered in this
study. TC’s in this work were determined using the second derivative approach using the
software by Tauxe (2002).
27
Table 1. Composition and Curie points (TC) of important ferromagnetic minerals for this
study (Dunlop and Özdemir, 1997).
Mineral
Composition TC (°C)
Magnetite
Fe3O4
580
Maghemite
Fe2O3
590-675
Hematite
Fe2O3
675
Titanomagnetite Fe3-xTixO4
150-540
Pyrrhotite
Fe7S8
320
Sometimes new magnetic minerals form by oxidation (or reduction) during the heating in
the laboratory (Dunlop and Özdemir, 1997). These are observed as irreversible heating
and cooling curves (for example sample ST64 in Fig. 16 of article III). Therefore the
temperature treatments show how chemical alteration induced by heating may affect the
magnetic behavior of samples. In Table 2 the possible chemical reactions for important
magnetic minerals are presented. Heating in argon gas may help to avoid some of these
changes.
Table 2. Alteration products of magnetic minerals during heating. The temperature at
which such process occurs is also given (Tarling, 1983; Dunlop and Özdemir,
1997).
Mineral
Alteration product T (°C)
Ti-magnetite Magnetite
> 300
Magnetite
Maghemite
150-200
Magnetite
Hematite
> 500
Maghemite
Hematite
350-450
Pyrrhotite
Magnetite
500
High temperature thermomagnetic (Curie-point) measurements have been performed on
specimens using the AGICO KLY-3S Kappabridge (articles I–IV). The powdered
specimens were heated from room temperature up to 700°C and cooled back to room
temperature in argon gas or in air, while susceptibility was measured during the whole
process (Fig. 8). This was done in order to determine the Curie temperature (TC, e.g. the
temperature above which a ferromagnetic mineral becomes paramagnetic) of a particular
magnetic mineral or phase.
Low temperature thermomagnetic measurements were also performed on selected
specimens (articles III and IV). To measure susceptibility at low temperatures, the
cryostat sleeve surrounding the specimen was filled with liquid nitrogen until the
specimen reached 80K. Liquid nitrogen was then drained off with argon gas before the
measurement series began. Specimen was heated up to 0 °C, while susceptibility was
28
monitored during the whole process (Fig. 8). This was done to determine the possible
transitions (e.g. Verwey for magnetite and/or Morin for hematite) of a particular magnetic
mineral or phase.
Figure 8. Example of high and low temperature thermomagentic measurement.
Thermomagnetic curves (susceptibility vs. temperature) for ABC samples of
the Vredefort impact structure. High temperature data shows a peak at ca.
580 °C typical for magnetite (i.e.Hopkinson peak). Dotted horizontal and
vertical lines indicate the change in the scale of the x-axis. In the first part of
the experiment specimens were heated (black line) from the temperature of the
liquid nitrogen up to 0°C. In the second part specimens were heated (black
line) from room temperature up to 700°C and cooled back to room
temperature (grey line).
2.4.2 Magnetic hysteresis properties
When a ferromagnetic material is magnetized using an external magnetic field, it will not
relax back to zero magnetization when the external magnetizing field is removed, hence
retaining some remanent magnetization. It must be driven back to zero by applying a field
in the opposite direction. A hysteresis loop shows the behavior of magnetization of a
particular mineral when it is cycled through a magnetic field.
A Princeton Measurements Co. Vibrating Sample Magnetometer (VSM) was used to
determine hysteresis loops at room temperature using a maximum field of 1 T. The small
chip or powdered rock sample undergoes sinusoidal motion by mechanical vibrations
within a uniform magnetic field. A magnetic material placed in a uniform magnetic field
will produce a dipole moment proportional to the product of the sample susceptibility and
the applied field. If the magnetic moment is made to undergo motion, the resulting
29
magnetic flux generates an electric voltage in the pick-up coil. This voltage signal will be
proportional to the magnetic moment, amplitude, and frequency of vibration.
The hysteresis loops provide many useful parameters to define the overall domain state of
a rock (Fig. 9). The saturation magnetization MS is the maximum magnetization achieved
by the specimen in the (maximum) applied field. The remanent saturation magnetization
MRS is the magnetization remaining in the specimen after the field is switched off. M S
and MRS strongly depend on the domain state (single domain (SD), multi-domain (MD)
or pseudo single domain (PSD)) of magnetic minerals present in the sample (Dunlop and
Özdemir, 1997). For example, typically magnetite saturates at 0.3 T, titanomagnetites
around 0.1 – 0.2 T, pyrrhotite at 0.5 – 1.0 T, and hematite around 1 – 6 T (O’Reilly,
1984; Dekkers, 1988). The coercive force HC is defined by the magnetic field needed to
reduce the magnetization to zero. The coercivity of remanence HCR is represented as the
field required to reduce the MRS to zero using the backfield technique (Dunlop and
Özdemir, 1997).
Figure 9. a) Example of hysteresis loop typical for randomly oriented SD or MD
magnetite grains. MS is the saturation magnetization (mAm2/kg); MRS is the
saturation remanent magnetization (mAm2/kg); HC is the coercive field (T). b)
Typical isothermal remanent magnetization (IRM) curve for randomly oriented
SD or MD magnetite grains. The saturation isothermal remanent magnetization
(SIRM) and coercivity of remanence (HCR) are shown. HSAT is the field needed
to saturate the magnetization.
For titanomagnetites, the ratios MRS/MS and HCR/HC as plotted in so-called Day-plot (i.e.
MRS/MS vs HCR/HC; Day et al., 1977) can be used to determine the granulometry (grain
size distribution) of the magnetic carriers in the specimens (for example Fig. 10). For
uniaxial single domain (SD) grains of magnetites the ratio MRS/MS is close to 0.5,
30
whereas for multi domain (MD) grains the ratio is close to 0.02. This difference resides in
the fact that the domain walls of MD grains are destroyed during the saturation. When the
field is removed, the walls do not return in their original position but settle in accordance
to imperfection in the lattice structure of the magnetic mineral. HCR/HC is less than 2 for
SD and higher than 5 for MD grains. The field in between SD and MD grain areas is the
area of pseudo single domain (PSD) grains. Such grains have a small number of domains,
and behave much as SD grains.
Figure 10. Measured hysteresis ratios for the Salla dyke plotted on the Day plot (Day et
al., 1977; Dunlop, 2002). Mr, Ms saturation remanence, saturation
magnetization; Hc coercive force; Hcr coercivity of remanence; SD single
domain; PSD pseudo single domain; MD multidomain.
2.4.3 Lowrie-Fuller- , Lowrie- and viscosity test
The domain structure of magnetic minerals and origin of the magnetization of
representatived specimens were investigated using a set of rock magnetic tests. The most
useful is the Lowrie–Fuller test (Lowrie and Fuller, 1971; see Article IV). This test is
suitable for titanomagnetites and is based on a comparison of the AF spectra of NRM, ,
ARM (ARM is used for approximation of TRM) and IRM (Fig. 11a). Dunlop (1983)
concluded, after trying to give a stronger experimental and theoretical basis for this test,
that the NRM and ARM behaviors are a manifestation of the coercivity spectrum which
is only indirectly related to domain size. First, the NRM acquired in general weak
31
geomagnetic field (40 – 60 PT) was stepwise AF demagnetized up to 160 mT (SQUID)
or 100 mT (LDA-3), depending on the instrument used. After that, an anhysteretic
remanent magnetization (ARM) was imparted with a peak AF of 100 mT and a biasing
field of 50 PT in the z-direction. The ARM was also demagnetized using the same
sequence. Finally, isothermal remanent magnetization (IRM) was induced in a 1.5 T field
along the z-direction and progressively AF demagnetized with the same scheme. To
discriminate between SD, MD, and PSD, the coercivity spectras of the NRM, ARM and
IRM are compared (Fig. 11a). In large MD grains, IRM requires larger destructive fields
than ARM to reach the same normalized remanence level. In SD grains, the opposite is
true. Following Dunlop (1983), the Lowrie-Fuller test can be quantized using parameter
MDF (median destructive field). MDF is calculated as the demagnetizing field required
halving the magnetization.
GH 1
MDF
2
H1
MDFNRM MDFIRM
MDFNRM
(4)
2
where H1/2 / H1/2 is change in the demagnetizing field needed to halve the magnetization,
MDFNRM is the demagnetizing field needed to halve the NRM (mT), and MDFIRM is the
demagnetizing field needed to halve the IRM (mT).
SD behavior is encountered in rocks with a grain size (or domain size) of the carriers
smaller than ca. 4 Pm, whereas the MD behavior corresponds to grains larger thanca. 15
Pm. The mixed behavior is defined by samples confined in the 4 – 15 Pm size interval.
The bimodal (BM) behavior represents samples having both SD and MD magnetic
carriers. Dunlop (1983) shows that MDF is negative for MD specimens. For SD or mixed
behavior it lies between 0 and 0.2, whereas in case of bimodal behavior it will be larger
than 0.3.
The Lowrie test (Lowrie, 1990) was used to utilize the different coercivity and thermal
unblocking temperature characteristic of the most common ferromagnetic minerals
(Article III). At first, three axis IRM were given in the following order: (1) 1.2 T in zdirection (diagnostic for hard magnetic carriers); (2) 0.4 T in y-direction (diagnostic for
intermediate magnetic carriers); and (3) 0.12 T in x-direction (diagnostic for soft
magnetic carriers) fields along the z, y and x axis, respectively. Then specimens were
thermally demagnetized in steps up to 680 °C. After each step magnetic susceptibility and
magnetization were measured. Finally, the magnetization of each orthogonal axis is
plotted against the temperature (Fig. 11b). The advantage of this method is that the
carriers of magnetization can be identified based on their coercivity and it may be
combined with the information of Curie temperatures.
32
Figure 11. a) Lowrie-Fuller test (Lowrie and Fuller, 1971) for pseudotachylitic (PT)
breccia sample (LV44) from the Vredefort impact structure. ARM of the sample
is harder than IRM, indicating that it contains single domain magnetic carrier.
b) Lowrie-test (Lowrie, 1990) for diabase sample (ST64) from Salla dyke. This
test indicates that medium to soft coercivity magnetite/titanomagnetite
dominates and pyrrhotite exists to a lesser extent.
Stability of the NRM was also studied using the simplified version of the so-called
viscosity test, which defines the changes in magnetization due to viscous processes (e.g.
Irving, 1964; Dunlop and Buchan, 1977; see Article III). It consists of measuring the
NRM of the specimens before (initial NRM = NRM0) and after (stable NRM = NRMst)
storing them in zero-field environment during three weeks. The viscosity coefficient Sv
(%) is defined:
SV
NRM 0 NRM st
u 100
NRM 0
(5)
where NRM0 is initial natural remanent magnetization (mA/m) and NRM st is the natural
remanent magnetization after storing samples in zero field environment (mA/m).
In general, SV values below 5% indicate that the NRM is stable.
2.4.4 Anisotropy of susceptibility
For reliable paleomagnetic studies the direction of the NRM should reflect accurately the
Earth’s ambient magnetic field (Irving, 1964). However, the NRM direction may deviate
from the prevailing field direction due to magnetic anisotropy of the rocks as a
remanence reorder. Anisotropy of the magnetic susceptibility may be caused by different
phenomena such as magma flow, tectonism or other biasing factors (e.g. Tarling and
33
Hrouda, 1993; Dragoni et al., 1997; Elming and Mattsson, 2001). In a rock specimen that
consists of randomly oriented magnetic grains the AMS would be close to zero, i.e. the
specimen is magnetically isotropic. However, if there is, for example a preferred
alignment of the magnetic grains due to tectonic forces, then the susceptibility takes the
maximal value along this axes (called as the maximum susceptibility axis). The two other
orthogonal axes refer to intermediate and minimal susceptibility values. It has been
suggested that the direction of remanence will not be significantly altered if the ratio
between the maximum and minimum susceptibilities (degree of anisotropy) is lower than
5% (e.g. Irving, 1964; Hrouda, 1982).
The anisotropy of magnetic susceptibility was measured for representative specimens
using the KLY-3 Kappabridge operating at 875 Hz frequency and 300 Am-1 (see Articles
II and III). The device allows automatic measurements of susceptibility on three
perpendicular axes while the specimen is spinning.
2.4.5 Optical methods – scanning electron microscopy studies
To verify the nature of the magnetic minerals, it is important to combine optical methods
with rock magnetic studies. In this study (Articles I, II, III and IV), we used a Jeol JSM5900LV scanning electron microscope of the Geological Survey of Finland to analyze the
mineralogy of representative specimens. Using optical methods, it is also possible to
estimate the alteration stage of the minerals.
2.5 Paleomagnetic direction and field tests
The direction of the paleomagnetic field is referred to the local horizontal plane, in terms
of a declination, D, which is the angle measured clockwise from true geographic north,
and an inclination, I, which is the angle measured up (regarded as negative) or down
(regarded as positive) from the horizontal plane. The magnitude and direction can be
illustrated on the same diagram as an orthogonal plot (e.g. Butler, 1992; for example of
plot see articles I – IV).
The basic requirement in paleomagnetic work is to verify that the observed ChRM is
original, recording a primary ambient field direction during emplacement of the rock
formation (or during the cooling of the impact melt sheet in case of impact structures) and
not during later events. Since in this study we deal mainly with ChRM of thermal origin it
can be tested by the baked contact (Fig. 12a; Everitt and Clegg, 1962) or the impact tests
(Fig. 12b; Pesonen, 2001). The concept of these tests is that an igneous body or impact
melt body 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 not likely to be greatly different from
the remanence direction observed from the formation itself. Due to the temperature
34
decrease the NRM directions of the baked zone gradually change. A complete test
requires (see Fig. 12): (1) an agreement in demonstrated stable remanence direction
between an igneous rock and adjacent baked country rock and (2) a difference between
this direction and the remanence direction of the unbaked host rocks. Otherwise there is a
possibility that both intrusion and the baked contact are later remagnetized (e.g. Irving,
1964; Butler, 1992). In article III we have shown a positive baked contact test, and in
Article I and IV we have indications of positive impact test.
Figure 12. a) Sketch of impact test (Pesonen, 2001). b) Sketch of baked contact test
(Everitt and Clegg, 1962). c) Example of a positive baked contact test carried
out for the samples from Salla diabase dyke (article III). Left: remanence
direcrtion of unbaked 2.2 Ga old gabbro. Middle: baked gabbro (2.2. Ga)
provided same direction than diabase samples. Right: remanence direction of
magnetization of 1.12 Ga old diabase sample. Arrows and letters indicate
direction of remanent magnetization.
35
Other field tests include conglomerate and fold tests. A positive conglomerate test
indicates that the characteristic remanence (ChRM) of the source rock has been stable at
least since the formation of the conglomerate. Butler (1992) summarizes that if ChRM in
clasts from a conglomerate has been stable since before deposition of the conglomerate,
ChRM directions from numerous cobbles or boulders should be randomly distributed
(positive conglomerate test). A non-random distribution indicates that ChRM was formed
after deposition of the conglomerate (negative conglomerate test).
In a fold test, relative timing of acquisition of a component of NRM (usually ChRM) and
folding can be evaluated (Butler, 1992). If a ChRM was acquired prior to folding,
directions of ChRM from sites on opposing limbs of a fold are dispersed when plotted in
geographic coordinates (in situ) but converge when the structural correction is made
(“restoring” the beds to horizontal). The ChRM directions are said to “pass the fold test”
if clustering increases through application of the structural correction or “fail the fold
test” if the ChRM directions become more scattered.
Stability of the paleomagnetic direction can be tested using the reversals test. It is based
on the fact that geomagnetic field changes polarity and that geocentric axial dipolar
nature of the geomagnetic field holds true during both normal- and reversed-polarity
intervals. At all locations, the time-averaged geomagnetic field directions during a
normal-polarity interval and during a reversed-polarity interval differ by 180°. ChRM
directions pass the reversals test if the mean direction computed from the normal-polarity
sites is statistically antiparallel to the mean direction for the reversed-polarity sites.
2.6 Paleomagnetic pole
The concept of the paleomagnetic pole was introduced by Creer et al. (1954) as a very
convenient method to present the paleomagnetic results (Fig. 13). The paleomagnetic
pole allows the paleomagnetic data from distant areas to be visualized in an easy and
practical way. First, it allows the time-successive data (variable ages) to be seen as
apparent polar wander path (APWP). Second, the deviation of two (or more) poles with
the same ages is crucial evidence that the sampling sites have been moving relative to
each other. This occurrence immediately calls for tectonic movements (e.g. tilting) or
movements of continents or lithospheric plates (blocks).
A pole position calculated, for example by spherical trigonometry, from a single
observation of the direction of the geomagnetic field (or from a single rock intrusion) is
called a virtual geomagnetic pole (VGP). This is the position of the pole of a geocentric
dipole that can account for the observed magnetic field direction at one location and at
one point in time. If VGPs are determined from many globally distributed observations of
the present geomagnetic field, these VGPs are scattered about the present geomagnetic
pole. A site-mean ChRM direction is a record of the past geomagnetic field direction at
the sampling site location during the (ideally short) interval of time over which the
ChRM was acquired. Thus, a pole position that is calculated from a single site-mean
36
ChRM direction is a VGP. As it is known that the Earth’s magnetic field changes polarity
and there are no adequate Precambrian apparent polar wander paths (APWPs) with
smooth transitions to Phanerozoic APWPs it is not possible to know a priori if the
studied Precambrian area was in the northern or in the southern hemisphere.
Figure 13. Determination of magnetic pole position from a magnetic field direction. Site
location is at S (s, s); site-mean magnetic field direction is Im, Dm; M is the geocentric
dipole that can account for the observed magnetic field direction; P is the magnetic pole
at (p, p); p is the magnetic colatitude (angular distance from S to P ); North Pole is the
north geographic pole; and E is the difference in longitude between the magnetic pole
and the site (Butler, 1992).
Consequently, most poles are now determined in the following manner: (1) from each
site-mean ChRM direction, a site-mean VGP is calculated; and then (2) the set of VGPs
is used to find the mean pole position by Fisher statistics (Fisher, 1953), treating each
VGP as a point on the unit sphere.
The Earth’s magnetic field undergoes secular variation (SV) which can be described by
changes in dipole (D) and nondipole (ND) components. The most important contribution
of SV to paleomagnetic data is due to westward drift of the ND field (Irving, 1964).
Effect of SV must be averaged out before paleomagnetic measurements are said to
conform with the GAD model. SV in paleomagnetic studies is expressed by the betweensite statistical scatter of the directions of the poles after the other causes for scatter, such
as the measurement errors or orientation errors (with-in site errors) are first eliminated
(e.g. Irving, 1962; McElhinny and McFadden, 2000). The GAD hypothesis states that, if
geomagnetic SV has been adequately sampled, the average position of the geomagnetic
pole coincides with the Earth's axis. Thus a set of paleomagnetic sites magnetized over
about 104 to 105 years should yield an average pole position (average of site-mean VGPs)
37
coinciding with the rotation axis. Pole positions calculated with these criteria satisfied are
called paleomagnetic poles (Butler, 1992). The difference between VGPs and the
paleomagnetic poles can be as much as 15-20° (e.g. McElhinny and McFadden, 2000).
In this study (one dyke, sill, or impact melt sheet), the only way to determine if SV has
been averaged out is to calculate the angular dispersion (between-site dispersion, S-value)
of the site-mean VGPs (e.g. Merrill and McElhinny, 1983) at the site paleolatitude. The
S-value is calculated from K (dispersion parameter for poles), and upper and lower limits
of S were calculated using the methods by Cox (1969):
S
6561
K
(6)
where K is the precision parameter from the pole position of each site.
In this work the angular dispersion is calculated and compared to the predicted dispersion
(e.g. model G by Merrill and McElhinny (1983); Fig. 16). If SV has been adequately
sampled, the observed angular dispersion of site-mean VGPs should be consistent with
that predicted for the paleolatitude of the sampling sites (however see also section 4.1.1
of this work). If the observed dispersion of site-mean VGPs is much less than predicted,
then the VGPs are more tightly clustered than expected for adequate sampling of secular
variation. A likely explanation is that the paleomagnetic sampling sites did not sample a
time interval covering the longer periodicities of SV (Butler, 1992). VGP dispersion
substantially larger than predicted indicates that there is a source of VGP dispersion in
addition to a sampling of SV. Perhaps there has been tectonic disturbance within the
sampling region or there is difficulty in determining the site-mean ChRM directions
(Butler, 1992).
In practice, even a collection of reliable paleomagnetic poles will have some scatter,
owing to imperfect averaging of geomagnetic SV, uncertainties in structural correction,
or other unknown effects.
2.7 Representing continental drift and building up the supercontinents
Based on a verified GAD hypothesis, it is known that the mean paleomagnetic pole
approximates the paleoposition of the rotation axis with respect to the continent from
which the paleomagnetic pole was determined. A paleogeographic position of a continent
can be created by rotating the pole together with the continent so that the paleomagnetic
pole is positioned on the axis of the geographic grid. The resulting map shows the
paleolatitude and the orientation (Fig. 14b) of the continent at the time when the magnetic
remanence was acquired. As the time-averaged geomagnetic field is symmetric about the
rotation axis, absolute values of paleolongitudes are arbitrary. Paleoclimatic indicators
such as evaporites (shallow mid-latitudes), coral reefs (equatorial latitues), glaciogenic
38
rocks (high latitudes), etc. see Irving (1964) are the best available independent measures
of paleolatitude with which to compare paleolatitudes determined from paleomagnetism
(e.g. Briden and Irving, 1964; Evans, 2006a). Latitudinal zones of climate exist
fundamentally because the flux of solar energy strongly depends on latitude. The present
mean annual temperature is 25°C at the equator but is only –25°C at the poles.
The technique of plotting an apparent polar wander path (APWP) for large paleomagnetic
data sets was introduced by Creer et al. (1954) and has become the standard method of
presenting paleomagnetic data. Fundamentally, an APWP is a plot of the sequential
positions of paleomagnetic poles from a particular continent, usually shown on the
present geographic grid (Fig. 14a). Through the GAD hypothesis, an APWP represents
the apparent motion of the rotation axis with respect to the continent of observation.
When APWPs were first developed, it was thought that APW was largely due to rotation
of the whole Earth with respect to the rotation axis, which is fixed with respect to the
stars. This whole-Earth rotation is known as true polar wander (TPW; e.g. Evans, 2003).
We now understand that for most of the last 200 million years (Besse and Courtillot,
2002), the major portion of APW was due to lithospheric plate motions carrying
continents over the Earth’s surface (continental drift; e.g. Butler, 1992). However, Evans
(2003) noted that there may be older periods for which TPW dominated between-plate
motions. There are also possible bursts of TPW within the last 200 million years
(Steinberger and Torsvik, 2008).
Figure 14. Left: the concept of an apparent polar wander path. Continent is held fixed
and paleopoles from t0 to t4 are connected to form an apparent polar wander
path. Right: the concept of continental drift. The same continent is shown in
its paleolatitudinal positions from t0 to t4, while holding the pole fixed
39
For testing the joint history of two or more continents during a specified time interval (for
example articles II and III), we compare if the APWPs of these continents match during
the proposed time interval (Fig. 15b). This principle is simply a corollary of the GAD
hypothesis for internally rigid plates. A paleomagnetic pole provides the past position of
the rotation axis with respect to the continent of observation. There can only be one
rotation axis at any particular geologic time. Thus, if two continents are placed in their
proper relative positions for a particular geologic time, their paleomagnetic poles for that
time must coincide. Furthermore, if these continents had the same relative position for a
significant interval of geologic time, their paleomagnetic poles (i.e. their APWPs) during
that entire time interval must coincide (Fig.15b). Considering that cratonic blocks for a
specific ancient tectonic association have not yet been established, Evans and Pisarevsky
(2008) documented examples of allowable cratonic conjunctions (or TPW) by comparing
great-circle distances between precisely coeval pairs of poles from each block. They also
found other examples by the same test that required differential motion between cratons
in Proterozoic times.
a)
b)
Figure 15. a) The principle of the closest approach technique for making reconstructions
(Pesonen et al., 2003a). A and B are hypothetical continents plotted in correct
latitudes and orientations as based on palaeomagnetic poles. Only northern
hemisphere option for continent A is shown. Positions a– d show continent B in
four possible positions along a latitude circle defined by palaeomagnetic data.
Positions c’ and c’’, for B, are allowed by the errors in palaeolatitudes.
Positions e– h in the opposite hemisphere (with B upside down) can also be
used in seeking better geological matchings between A and B. b) Testing if the
APWPs of two now separate continents match during the proposed time
interval. Left: Observed APWPs. Right: Rotation of observed APWPs using
same Euler pole produce matching APWP (modified from Pesonen and Sohn, in
press).
40
Precambrain paleopoles are problematic, as there are commonly large uncertainties
concerning their ages. This makes it impossible to properly sequence them along APWPs.
Long time gaps in the paleopole record from a given continental block make it difficult to
interpolate between poles. In addition, interpolation between widely separated paleopoles
is hindered by an ambiguity in magnetic polarity that results from the lack of continuous
APWPs from the Precambrian to the present. Buchan et al. (2000; 2001) suggested the
direct comparison of individual key paleopoles of the same age from different continental
blocks to determine relative positions of these blocks. This would not yield a unique
reconstruction, because of the longitudinal uncertainty inherent in the paleomagnetic
method applied to individual poles (Fig. 15a).
In addition, Pesonen et al. (2003a) list paleomagnetic criteria in validating the
reconstructions: (1) the magnetic polarity had to be the same from one continent to
another during the lifetime of the assembly; and (2) the drift rates and the sense of drifts
of the cratons had to be comparable during the lifetime of the assembly. In addition to
paleomagnetism (including study of magnetic anisotropy, polarity, APWP, intensity), one
needs to take into account certain geological indicators for testing supercontinent, such as
dyke swarms, age patterns, orogenic belts, and anorogenic rocks to reconstruct ancient
positions of continents and build up proposals of paleogeography of supercontinent
assemblies. Paleoclimatology (latitudinal indicators) and paleobiogeography (e.g. fossils)
may give further support for reconstructions.
3. Summary of the articles I - IV
3.1 Article I
Salminen, J., Donadini, F., Pesonen, L.J., Masaitis, V.K., and Naumov, M., 2006.
Paleomagnetism and petrophysics of the Jänisjärvi impact structure, Russian Karelia.
Meteoritics and Planetary Science, 41, 1853-1870.
In this study paleomagnetic, rock magnetic and petrophysical properties of the target and
impact rocks of the Lake Jänisjärvi meteorite impact structure were investigated for three
reasons. First, to achieve additional data for the apparent polar wander path (APWP) of
Baltica for the age interval of 750-650 Ma, which is poorly defined. Second, the
Jänisjärvi data have potential not only to define the position of Baltica at ca. 750-650 Ma
but also to test whether supercontinent Rodinia was already fragmented (e.g. Li et al.,
2008) and to reconstruct the post-Rodinia paleolocation of Baltica. Third, we tried to
determine the ancient geomagnetic field intensity using impact melted rocks in
paleointensity experiments.
The Lake Jänisjärvi impact structure (latitude = 61q58’N, longitude = 30q58’E) is located
in Russian Karelia, about 220 km north of St. Petersburg. Based on gravity data the
diameter of the structure is about 16 km. The structure is located close to the boundary
41
between Archean and Proterozoic terrains of the Baltic Shield. Impact rocks (melt,
suevites, and breccias) outcrop in central islands off cape Leppiniemi on the western part
of the lake. The Jänisjärvi structure is surrounded by Proterozoic crystalline schists,
which are strongly fractured due to the impact event. The age of the impact melt
(tagamite) samples have been dated to 718 ± 5 Ma (K-Ar method; Masaitis et al, 1976)
and 698 ± 22 Ma (39Ar-40Ar method; Müller et al., 1990). More precise 39Ar-40Ar data
(682 ± 4 Ma Jourdan et al., 2008) appeared subsequent to our paper's publication, and the
slightly revised conclusions will be discussed in section 4.1.2 of this thesis.
A characteristic remanence component (ChRM) D = 101.5°, I = 73.1° (95 = 6.2°) was
observed from impact melt specimens (155 specimens out of 22 hand samples) carrying
stable remanence. Such a component is interpreted to be related to the Jänisjärvi impact
event and can be partly observed as an overprint in target rocks thus yielding a positive
impact test (Pesonen, 2001). The ChRM of the impact melt specimens yields a pole
position of Plat = 45.0°N, Plon = 76.9°E (A95 = 10.4°), which places Baltica on latitude
of 60° at the time of acquisition of ChRM.
Data for Table 3 presented in article I was revisited for verification of the positive impact
test. The reversed directions for component C (impact overprint) as listed in Table 3 in
article I have to be negated being too shallow. However, samples AL1, JH1, JH2, JC1,
JC2, JC3 and JD1 from near the rim of the structure indeed show an impactite type
component (component C) as an overprint. Samples JH1, JC1, JC2, JC3 and JD1 show
both impactite type component (component C) and Svecofennian component (component
A). Samples farther from the rim of the structure (for example JH4 and JH13) show only
component A (Svecofennian component). Moreover, there is further evidence of a
Svecofennian component (component A) in various dykes on Russian Karelia (e.g.
Mertanen et al., 1999; Mertanen et al., 2006) as well as in the eastern Finland not far from
Jänisjärvi (Pesonen et al., 1991). Therefeore, the paleomagnetic data within and beyond
the Jänisjärvi structure support a positive impact test (Pesonen, 2001).
Paleointensity studies of melt samples show remarkably higher paleointensities compared
to other paleointensity results of Neoproterozoic-Cambrian rocks, although the
paleointensity data has a gap in data about this age (e.g. Donadini, 2007)..
The comparison of the Jänisjärvi pole on the APWP of Baltica shows disagreement with
published isotopic ages and observed pole positions. This discrepancy may be due to
inherent problems of the 39Ar-40Ar dating technique or because of several problems in
paleomagnetic data: (1) post-impact remagnetization; (2) post-impact tilting of the
sampling sites; (3) secular variation of the geomagnetic field, which is not averaged out;
and (4) a poorly defined APWP due to the paucity of well-dated Neoproterozoic
paleomagnetic poles.
According to Deutsch and Schärer (1994), the whole rock K-Ar and 39Ar-40Ar ages of
impact melt samples might be ambiguous for the following reasons: (1) the post-shock
loss of in situ produced radiogenic argon (40Arrad) by devitrification and diffusional
processes can result in ages that are too young; (2) the presence of relic 40 Arrad in the melt
42
matrix, as well as in rock and mineral fragments, which are inherited from precursor
lithologies, can cause ages that are too old; and (3) disturbances by secondary processes,
such as incorporation of Ar from a fluid phase can lead to either increase or decrease of
the age. [See section 4.1.2 in this work for further discussions]
The positive impact test rules out the possibility of the post-impact remagnetization. The
well documented 0.2° tilting of the pre-Vendian peneplain of the Fennoscandian Shield
(e.g. Laitakari et al., 1996) does not affect the pole position. Moreover, as the impact melt
layer appears on the same stratigraphic level everywhere in the Jänisjärvi structure
(Masaitis, 1999) and local tilting has not been reported, the secular variation (SV) was
investigated using the method by Pesonen et al. (1992). It gave hints that the SV may not
have been fully averaged out in which case the pole may be a virtual geomagnetic pole
[see also section 4.1.2 in this work]. Also, because we are studying a single impact melt
layer, it is most probable that SV has not been averaged out. A large gap in APWP of
Baltica from 750 to 616 Ma makes the comparison of Jänisjärvi pole with the APWP
problematic, as it involves interpolation of the APWP at 700 Ma.
Our observations support the quality and primary origin of the Jänisjärvi paleomagnetic
pole, although it is a virtual geomagnetic pole [see section 4 of this work for futher
discussions].
3.2 Article II
Salminen, J., and Pesonen, L.J., 2007. Paleomagnetic and rock magnetic study of the
Mesoproterozoic sill, Valaam island, Russian Karelia. Precambrian Research, 159, 212230.
This research concentrates on the 1458+4/-3 Ma monzodioritic sill of the Valaam island,
Lake Ladoga, Karelia. The sill includes intruded syenitic plugs, dykes, and veins. The sill
is part of the mid-Proterozoic (1.67–1.46 Ga) anorogenic magmatic activity, which
formed the magmatic province that extends west from Russian Karelia towards southern
Finland, central Sweden (e.g. Tuna dykes in Sweden with an age of 1.46 Ga), and south
to the Baltic countries where they are overlain by younger sediments (e.g. Söderlund et
al., 2005; Rämö and Haapala, 2005; Luttinen and Kosunen, 2006).
The study was carried out in order to shed light into the configuration and life-time of the
supercontinent Hudsonland (Columbia, Nuna) and in particular, focusing on the question
of the proposed long-lasting unity of Baltica and Laurentia (from 1.83 to 1.26 Ga;
Pesonen et al., 2003a). The Valaam sill provides important material for a paleomagnetic
study for two reasons: (1) its magnetization is stable, which was shown by a preliminary
study of Pesonen (1998) and appears to be distinct from the directions of Jotnian diabases
(1.26 Ga) and Subjotnian units (1.67 – 1.5 Ga); (2) there is a lack of well dated poles
between 1.6 and 1.3 Ga for Baltica (e.g. Buchan et al., 2000; Pesonen et al., 2003a).
43
Th epaleomagnetic study demonstrated the presence of two remanence components: a
well defined high coercivity characteristic remanent component (ChRM) (D = 43.4°, I = 14.3°, D95 = 3.3°) and a viscous component interpreted to be related to the present Earth’s
magnetic field. According to thermomagnetic, high field and scanning electron
microscope analyses, the ChRM is carried by pseudo single domain or multi domain
titanomagnetite grains with varying Ti-content. Although the baked contact or fold tests
were not possible, the paleomagnetic data of the sill, of other units around Lake Ladoga
(similar ages to Valaam sill), and those of Svecofennian rocks (age of ca. 1.9-1.8 Ga)
around lake Jänisjärvi (article I), suggest a primary origin for the Valaam ChRM.
Calculated cooling time of the Valaam sill is with thickness of 130-150 m (assumed
depth 2 km) is order of hundreds of years (430-600 years using the method of Jaeger
(1957) or 200 years using the method of Delaney (1981)), much less than time needed to
average out the secular variation. The Valaam ChRM yields a virtual geomagnetic pole
Plat = 13.8°, Plon = 166.4° (A95 = 2.4°), which places Baltica at low paleolatitude (ca.
7°). This is also supported by the paleoclimatology, as evidenced by Mesoproterozoic ca.
1.5-1.3 Ga red-bed sedimentation (e.g. the Satakunta sandstone; Rämö et al., 2005). As a
main result of this study a new supercontinent Hudsonland configuration for ca. 1.46 Ga
is proposed. According to combined result from Lake Ladoga area and Tuna dykes
northern Baltica is placed against north-eastern Greenland. This differs for example from
Meert (2002) and Pesonen et al. (2003a) models, but is similar to those supported by
geological data (e.g.Gower et al., 1990; Gorbachev and Bogdanova, 1993). Further
support for this reconstruction is the continuity of the rapakivi magmatism from Baltica
to Laurentia at 1.6-1.5 Ga (Rämö et al., 2005). According to this study the existence of a
single continental Baltica-Laurentia block since 1.76 Ga until 1.27 Ga is
paleomagnetically permissible and makes us consider the possibility that the Hudsonland
configuration lasted such a long time. [For updated discussion and pole definition see the
Section 4.1 of this work]
3.3 Article III
Salminen, J., Pesonen, L.J., Mertanen, S., Vuollo, J., and Airo, M.-L., 2009.
Palaeomagnetism of the Salla Diabase Dyke, northeastern Finland and its implication to
the Baltica – Laurentia entity during the Mesoproterozoic. Manuscript to be published in:
Reddy, S.M., Mazumder, R., and Evans, D.A.D., (eds.) Palaeoproterozoic
Supercontinents and Global Evolution. Geological Society of London, Special
Publication, 323.
It has been proposed (e.g. Pesonen et al., 2003a) that after the breakup of supercontinent
Hudsonland Baltica separated from Laurentia at ca. 1.25 Ga and started its journey
towards the southern latitudes. Baltica, Laurentia and most other continents are known to
reassemble again at ~1.2-1.1 Ga to form the supercontinent Rodinia. The paleoposition of
Baltica at ca. 1.2-1.1 Ga ago is of fundamental importance to precisely define the
amalgamation period of Rodinia and to study the proposed long-lasting unity (see Article
44
II) of Baltica and Laurentia forming the core of Hudsonland. To accomplish this task, we
have carried out paleomagnetic and rock magnetic studies on the well dated 1122 r 7 Ma
(U-Pb, Lauerma, 1995) Salla diabase dyke in northern Finland. This dyke, being 60-140
m wide and more than 130 km long, cuts the older Paleoproterozoic rocks and constitutes
the youngest rock formation in north-central Finland. This magmatic activity at ca. 1.12
Ga is a phenomenon as coeval rocks are available in Laurentia (oldest Keweenawan
units), Australia (Lake View dolerites), and elsewhere.
Altogether 194 diabase samples from 15 sites were collected along the dyke. In order to
carry out the baked contact tests, samples were also taken from the country rocks. The
results of the paleomagnetic study show that the Salla diabase dyke yields a stable
characteristic remanence component (ChRM) and a viscous overprint. The direction of
the ChRM (D = 42.2°, I = 73.9°, 95 = 4.8°) is rather close to the direction of the present
Earth’s magnetic field at the sampling site. However, a positive baked contact test with
2.2 Ga gabbro dyke (remagnetized at 1.9 Ga?) proves the primary nature the ChRM.
Rock magnetic and scanning electron microscope studies indicate that (titano)magnetite
is the carrier of ChRM.
Unfortunately, although the distance between sites is more that 90 km, the sites come
from the single huge dyke and therefore, secular variation may not have been fully
averaged out. In spite of this the new virtual geomagnetic pole (VGP; see also section
4.1.1 of this work, Fig. 17) provides an important result to define the late
Mesoproterozoic position of Baltica. The VGP of Plat = 71°N, Plon = 113°E (A95 = 8.1°)
is distant from any known Proterozoic paleopole of Baltica and therefore, the preSveconorwegian apparent polar wander path (APWP) of Baltica must be modified. As the
the pole is a VGP, which may differ from the true 1.12 Ga paleopole by 15-20q (e.g.
McElhinny and McFadden, 2000) plus its own A95 (8.1q), the suggested loop may be
anything between ca. 50q and ca.100q. A nearly coeval Logan Loop (e.g. Robertson and
Fahrig, 1971; Ernst and Buchan, 1993; Donadini, 2007) of the Laurentian APWP and
1165 – 1000 Ma APWP of Kalahari-Grunehogna craton also express large loop-like
APWPs, which may imply a common drift history for these continents (see Pesonen et
al., 2003a). However, as Evans (2003) emphasized, this could be due to the cratons being
part of a supercontinent or that the TPW is dominating the APWP during that interval.
New paleomagnetic data from well dated rocks between 1270 – 1112 Ma are required
from all of these continents to test the similarity between these APWP loops.
Nonetheless, further evidence of the common history of the Baltica and Laurentia is the
fact that the similarity between the APWPs begins already at 1.75 Ga.
Extension of the proposed long-lived Mesoproterozoic connection between Baltica and
Laurentia (e.g. Gower et al., 1990; Åhäll and Connelly 1998; Karlstrom et al., 2001;
Pesonen et al., 2003a and references therein) is permissible according to the new 1.12 Ga
paleomagentic data of the Salla dyke supporting a long-lived (1.76 – 1.12 Ga) Baltica –
Laurentia connection. In this pre-Rodinia (ca. 1.12 Ga) configuration of Baltica and
Laurentia, Baltica is 90° rotated compared to Rodinia models of Hoffman (1991),
Pesonen et al. (2003) and Li et al. (2008), but being consistent with configuration of
Evans (2006a). The new VGP of Salla dyke provides new constraint on the timing of the
45
rotation of Baltica relative to Laurentia (e.g. Gower et al., 1990). However, this
interpretation needs further verification when taking into account that the ages of the
compared poles for Baltica and Laurentia are not exactly coeval and the data for Baltica
comes from only one dyke where secular variation has not been adequately averaged.
3.4. Article IV
Salminen, J., Pesonen, L.J., Reimold, W.U., Donadini, F., and Gibson, R.L., 2008.
Paleomagnetic and rock magnetic study of the Vredefort impact structure and the
Johannesburg Dome, Kaapvaal Craton, South Africa – implications for the apparent polar
wander path of the Kaapvaal craton during the Mesoproterozoic. Precambrian Research,
168, 167-184.
New paleomagnetic and rock magnetic results are presented for various lithologies from
the region of the Vredefort impact structure (2023 ± 4 Ma (2); Kamo et al., 1996) and
from the Johannesburg dome, both located in the Kaapvaal Craton, South Africa. The
Vredefort impact structure is the oldest, well-dated, and with an estimated original
diameter of 250–300 km (Henkel and Reimold, 1998), the largest impact structure
identified on Earth (for a review of the Vredefort impact structure see Gibson and
Reimold, 2001, 2008). This study was conducted because rocks from the Kaapvaal
Craton are quality candidates for providing paleomagnetic data in the period between the
time interval for the Kenorland supercontinent (2.45–2.1 Ga) and formation of the Early
Proterozoic Hudsonland supercontinent (1.83 to 1.50–1.25 Ga).
However, earlier paleomagnetic studies of Vredefort lithologies have revealed the
problematic nature of the Archean rock within the Vredefort impact structure, namely the
extraordinarily high values of natural remanent magnetization (NRM) interpreted to be
caused by the impact event by various mechanisms (Hart et al., 1995; 2000; Pesonen et
al., 2002; Carporzen et al., 2005). A detailed paleomagnetic and rock magnetic study was
undertaken of samples from the Vredefort impact structure. The sampling region was
extended to the Johannesburg Dome to further elucidate these problematic issues and to
attempt to evaluate the paleomagnetic potential of this region of Kalahari for global
paleoreconstructions.
A series of pre-impact alkaline and mafic intrusive complexes occur in and around the
outer reaches of the Vredefort Dome. With this paleomagnetic study, we also contribute
the discussion raised by Hargraves (1970; 1987) concerning the emplacement timing of
the Rietfontein, Roodekraal, Schurwedraai, and Lindeques Drift intrusions relative to the
impact-generated overturning of the supracrustal strata occurring in the outer collar parts
of the Vredefort Dome. Yet, as the data for samples from the Rietfontein, Schurwedraai,
Lideques Drift, and Roodekraal intrusions are quite sparse, of low quality, and quite
scattered, we are unable to prove or disprove the collar overturning.
46
We also investigated whether the samples from various lithologies show any overprinting
from younger geological events such as the emplacement of the Anna’s Rust Sheet (ARS)
gabbro (ca.1.1 Ga), events related to the emplacement of the Umkondo Large Igneous
Province at ca. 1.1 Ga, the accretion of the Namaqua-Natal Belt (ca. 1.1 Ga) onto the
Kaapvaal Craton, or the emplacement of Karoo basalts (0.185 Ga).
Finally, this study also contributes to the controversial discussion regarding suggested
temperatures in the core of the Vredefort structure before and after the impact event
(Carporzen et al., 2006; Gibson and Reimold, 2005; Gibson et al., 1997; Henkel and
Reimold, 2002; Turtle et al., 2003; Muundjua et al., 2007).
The first important result from this study is the discovery of high Q values in samples not
only in all pre-impact lithologies but also in some Vredefort impactite samples, in the
much younger (1.1 Ga) ARS samples and even in the Johannesburg Dome samples (ca.
120 km away from the Vredefort structure). As the high Q values are also observed in the
Johannesburg Dome sample, a direct link to the Vredefort impact as a cause can be ruled
out. As the observed NRM is rather hard and shows multiple components, we exclude
lightning as a cause for all observed high Q values (except in the case of ARS gabbro
samples, which have high Q values and are magnetically very soft). Instead, the cause of
the high Q values may be related to fluid circulation within the two domes and the high
temperatures of the rocks which were uplifted by the impact event from a mid-crustal
original setting. Exposing high temperatures and hot fluids made the rocks vulnerable to
acquire high thermochemical remanence. The uplift (“doming”) resulting from elevated
temperatures can explain the enhanced Q values in both dome areas (Vredefort,
Johannesburg), but the role of impact is still a puzzle. [Further discussions in section 4.2
of this work].
Paleomagnetic research of Vredefort impactites reveal a stable, hard coercivity and high
blocking temperature characteristic remanence component (ChRM) with D = 18.3°, I =
54.8° (95 = 8.1°) (12 sites, 16 samples, 56 specimens), which is carried by magnetite.
Such a direction indicates a pole position of Plat = 25.1°N, Plon = 43.5°E (A95 = 10.6°),
which places the Kaapvaal Craton at a paleolatitude of 35° at ca. 2.02 Ga. Owing to the
small number of samples that yielded sufficiently meaningful results, it is possible that
secular variation has not been sufficiently averaged out. In that case, the pole represents
only a virtual geomagnetic pole, indicating that its use for paleoreconstructions may be
speculative. However, the observed direction of magnetization is agreeable with results
by Hargraves (1970), Hart et al. (1995; 2000), and Carporzen et al. (2005). Also, because
such a ChRM component was isolated as well from the Archean basement rocks within
the Vredefort Dome, we consider this “impact” component to be primary. The pole falls
onto the Paleoproterozoic (2200-1900 Ma) swathe of the apparent polar wander path of
the Kaapvaal Craton.
Paleomagnetic analysis of the rocks around the 1.1 Ga ARS gabbro intrusion in the
northern part of the Vredefort structure revealed the presence of either a shallow north or
a shallow south direction, which is tentatively related to emplacement of dyke, sills, and
lavas of the Umkondo Large Igneous Province. The analysis of all rocks, including the
47
Vredefort impactites, yields rare, distinct great circle paths towards these shallow
directions as also observed by Jackson (1982) and Layer (1986). A likely explanation for
this overprint direction is the heating and hot fluid activity caused by now-eroded ARStype gabbros in the area. No evidence of a Karoo-type (0.18 Ga) overprint is seen in
Vredefort impact rocks.
Low temperature vs. susceptibility measurements (Verwey transitions) indicate that the
distinction of pre- and syn-impact magnetite grains is not as straightforward as proposed
by Carporzen et al. (2006). In addition, some Archean basement samples from the central
part of the Vredefort dome with typical NRM intensity values show an impact-related
magnetization component carried by magnetite, which supports the idea that the impact
structure experienced temperatures needed to reset the magnetization direction of
magnetite (~580 °C). Moreover, textural evidence indicates that immediately after the
passage of the impact-generated shock wave, the maximum temperatures experienced by
rocks diminishes radially away from the center towards the periphery of the structure of
the Vredefort Dome (Gibson and Reimold, 2005; Gibson et al., 1997). There is
mineralogical evidence that the temperatures of the rocks currently exposed in the central
core (radius 8 km) of the Vredefort Dome exceeded the minimum granite solidus
(~650–700 °C; Gibson, 2002; Gibson et al., 1997).
3.5 Author’s contribution
The present author’s own contribution to each publication is shown in Table 3, classified
into two categories: (1) technique (includes collection of samples, preparation, and
measurements); and (2) processing (includes data analysis, interpretation, and writing of
the articles).
Table 3. Percentages of the present author’s own contribution to each publication.
Article
I Jänisjärvi impact structure
II Valaam sill
III Salla dyke
IV Vredefort impact
Technique
40
100
70
60
Processing
85
>95
>95
>95
In Article I (the Jänisjärvi impact sructure), the first author collected and prepared 10% of
the samples (together with L.J. Pesonen and F. Donadini) and processed all the data
except the paleointensity measurements, which is based on work by F. Donadini
(Donadini, 2007). The first author wrote the majority of the article except for the section
regarding paleointensity. Donadini is credited for Table 4 and Figures 1, 8, and 11. The
general paleomagnetic background behind the study was based on the MSc work of the
first author. A detailed rock magnetic study was carried out over the course of this
dissertation work.
48
In Article II (the Valaam sill), the first author collected the samples during field
campaigns in 2003 (with L.J. Pesonen and F. Donadini) and 2004 (with L.J. Pesonen and
T. Kohout). The first author also prepared the samples, made all measurements, and was
responsible for the data analysis, interpretation, and publication.
In Article III (the Salla dyke), the first author collected 30% of the samples (with L.J.
Pesonen, S. Mertanen, J. Vuollo, and S. Raiskila), prepared, and measured them. The
author also measured additional specimens from an existing collection of samples, except
one set collected by S. Mertanen. The author made rock magnetic measurements,
analyzed data, and wrote the article. The other authors are credited for Figures 1 and 2.
In article IV (the Vredefort impact structure), the present author collected 20% of the
samples (with U. Reimold and R.L. Gibson), prepared, and measured all of them. She
also made additional paleomagnetic measurements of samples from existing collections
by U. Reimold, R.L. Gibson, and L.J. Pesonen. Also, she is responsible for the majority
of rock magnetic measurements and data analysis, interpretations, and the majority of the
writing. The other authors are responsible for Figures 1, 2, and 4e.
4. Updated discussion and synthesis of the results
4.1 New paleomagnetic results for Baltica
The main objective of this study was to provide new poles for the apparent polar wander
path (APWP) of Baltica. Observed virtual geomagnetic poles from Valaam monzodioritic
sill (article II), from Salla diabase dyke (article III) and from Jänisjärvi impact melt
(article I) are listed in Table 4.
Table 4. Obtained new paleomagnetic data for Baltica.
Rock unit
Valaam monzodioritic sill
Salla diabase dyke
Jänisjärvi impact melt
Age
(Ma)
1458+4/-33
1122±72
682±41
B/N/n
9*/36/154
13*/170/484
5*/22/155
D
(°)
43
42
102
I
(°)
-14
73
73
95
(°)
3.3
4.8
6.2
Plat
(°)
14
71
45
Plon
(°)
166
113
77
A95
(°)
2.4
8.1
10.4
1
Jourdan et al., 2008 (see also discussion below). 2 Lauerma, 1995. 3 Rämö et al., 2001; 2005.
B/N/n are number of sites/samples/specimens. * denotes the statistical level used for mean
calculations. D and I are the declination and inclination of the remanent magnetization
component. D95 is the radius of circle of 95% confidence of the mean directions. Plat and Plon
are the latitude and longitude for the pole. A95 is the radius of the circle of 95% confidence of
the pole. S is the estimated angular standard deviation of the pole.
49
S
(°)
12.8
15.6
18.2
4.1.1 Paleomagnetic results from the Valaam sill and from the Salla diabase dyke – the
proposed long-lasting connection between Baltica and Laurentia during the
Mesoproterozoic
The ChRM of Valaam sill samples (1458+4/-3Ma, Rämö et al., 2005) are well defined
(Table 4) and similar in both syenitic and monzodioritic samples, indicating that the
samples recorded the ambient field simultaneously. Moreoever, the shape of thermal
demagnetization curve of monzodioritic samples is “a shoulder type”, indicating that the
magnetization is not an overprint (e.g. Dunlop and Özdemir, 1997). Although baked
contact or fold tests were not possible, the paleomagnetic data of the Valaam sill and of
the other coeval units, such as Tuna dykes (Bylund, 1985) and Early Riphean dyke
complex from the Lake Ladogan area (Shcherbakova et al., 2008), are similar. Whereas
the paleomagnetic data of the older Svecofennian (ca. 1.9 Ga) rocks around Lake
Jänisjärvi, some 40 km North of Lake Ladoga, are very different (see section 4.1.2 of this
work; Mertanen et al., 1999 and reference therein). This suggests a primary origin for the
observed ChRM component of the Valaam sill. When comparing the observed betweensite dispersion (S) of site-mean paleopoles for this component we obtain a value of 12.8°
(Table 4 and Fig. 16) which is consistent with the paleolatitude of Baltica (7°) suggesting
that the geomagnetic secular variation (SV) has been adequately sampled. However,
concordance of dispersion with a geodynamo SV model is no guarantee of adequate
averaging, as error can result from a number of alternative effects. Since, Valaam sill is a
single intrusive body (estimated thickness of 130–150m; Amantov et al., 1996) and
calculated cooling of the sill (in depth of 2 km from intial temperature of magma 11501100 °C below 580 °C) is order of hundreds of years (ca. 200 years using the method of
Delaney, 1981 and ca. 430-600 years using the method of Jaeger, 1957), much less than
required to average secular variation out (order of 105 years). Therefore, the “matching”
of the S-parameter with the standard (G-model) SV curve may be fortuitous. Thus, the
pole must be treated as a virtual geomagnetic pole (VGP).
Figure 16. Calculated between site dispersion (S) of observed poles for Baltica are
plotted on the secular variation Model G of Merrill and McElhinny (1983).
50
Table 5. Combined ca. 1460Ma paleomagnetic poles for Baltica.
Rock unit
Age
(Ma)
Plat
(°)
Plon
(°)
A95
(°)
1458+4/-3
14
166
2.4
1460
15.4
178
3.2
1460
9.9
174
3.6
1460
4.7
167.2
4.8
Riekkalansaari
1460
15.1
181.1
4.4
MEAN
1460
11.9
173.2
7.5
Valaam monzodioritic sill
1
Tamkhanka norher outskirts
Tamkhanka southern outskirts
Suur-Haapasaari1
1
1
1
Shcherbakova et al., 2008. Plat and Plon are the latitude and longitude for the pole. A95
is the radius of the circle of 95% confidence of the pole.
We combined the result from the Valaam sill with selected ca. 1460 Ma well-dated units
of Baltica (Fennoscandia, Lake Ladogan area) (see Table 5). Further results – not
included in Table 5 – from recently well-dated (1461-1462 Ma; Söderlund et al., 2005)
dolerite dykes (Gallsön, Bunkris, Glysjön, Tuna and associated Gustaf porphyries)
located in SW Sweden are forthcoming (Lubnina et al., 2007). For example, existing
paleomagnetic data by Bylund (1985) from the Bunkris-Glysjon dyke is not included.
The paleomagnetic results from this dyke (as discussed by Söderlund et al., 2005) vary
along its strike (Bylund, 1985; Bylund and Elming, 1992) and in a remarkably consistent
manner among all studies. Bylund and Elming (1992) report data from the Gallsjön dyke
(1461.2±1.4 Ma; Söderlund et al., 2005). The Gallsjön direction of magnetization is the
same as that of the much younger Falun dyke and similar to the 935-Ma Göteborg dyke
pole of Pisarevsky and Bylund (2006) indicating that it might be overprinted. Finally, the
Bunkris-Glysjon dyke may feed the Öje basalt, but this data (Piper and Smith, 1980) is
not included in the grand mean pole for 1460 Ma. Also, we are able to obtain a robust
paleomagnetic pole of Plat = 11.9°, Plon = 173° (A95 = 7.5°) (Table 5) from formations in
Karelia, which places Baltica either on the southern hemisphere (10°S), or geographically
inverted in the northern hemisphere (10°N). Further support for the low-latitude location
of Fennoscandia comes from the paleoclimatological data, which includes the
Mesoproterozoic ~1.5-1.3 Ga red-bed sedimentation (for example the Satakunta
sandstone; Rämö et al., 2005).
The samples from the Salla diabase dyke (1122 ± 7 Ma; Lauerma, 1995) provided a
stable ChRM component (D = 42 °, I = 73 ° and 95 = 4.8°). The 95%- error circle of this
ChRM component encloses the direction of the Present Earth’s field (D = 11°, I = 76°) on
the sampling site. However, on the basis of a positive baked-contact test to the 2.2 Ga
gabbro sill at site Jokinenä, the fact that the diabase carries clearly two components
(viscous and stable ChRM), and a viscosity test which indicates stable remanent
magnetization, we suggest that the origin of the ChRM component is thermal and
represents the primary remanence of the diabase dyke acquired during the cooling of the
dyke. Altough we collected 13 separate sites over a distance of ca. 90 km the results
represent only one dyke. Since, thickness of the is 60–100 m; Amantov et al., 1996) and
calculated cooling of the center part of the dyke (from intial temperature of magma 115051
1100 °C below 580 °C) is order of hundreds of years (ca. 40-100 years using the method
of Delaney, 1981 and ca. 90-270 years using the method of Jaeger, 1957), much less than
required to average secular variation out (order of 105 years). Hence it is possible that the
geomagnetic secular variation has not been sufficiently averaged out and the pole
represents only a virtual geomagnetic pole (VGP). Even though the upper limit of the Sparameter (Table 4, Fig. 16) falls on the model G curve of Merrill and McElhinny (1983),
the mean value (15.6; Table 4) falls below the curve suggesting that the observed pole
may be a VGP. Nonetheless, tentative evidence for the positive reversals test comes from
seven rather stable samples taken from the Haltiavaara site (e.g. Fig. 17), which yield a
mean direction of D = 349°, I = -86° (k=38°, 95= 7°, 7 samples, 15 specimens) not
documented in article III. However, these results must be confirmed. In addition, the
direction is non-antipodal due to a rather steep inclination. The presence of seven
reversed polarity samples near the contact part of the dyke suggests that the cooling of
this huge (thick) dyke may have lasted longer than several hundreds of years, maybe even
thousands of years. This is peculiar since cooling of the center of the dyke with width of
60-100 m from 1150-1000 °C below 580 °C should not take more than some hundreds of
years as discussed above. However, dyke intrusion could set up a hydrothermal system
along its margin, that persisted for much longer than the timescale of conductive cooling
(e.g. Delaney, 1987).
Figure 17. Observed reverse direction of magnetization in diabase sample ST4 taken
from site Haltiavaara, Salla diabase dyke.
The pole (Plat = 71°N, Plon = 113°E, and A95 = 8.1°) of the Salla diabase dyke is not
close to any previously known Proterozoic poles of Baltica. It adds a clockwise loop to
the APWP of Baltica between 1265 and 1040 Ma, just before the Sveconorwegian Loop
(Fig. 18). Since it is a VGP, which may differ from a time-averaged paleopole by 15-20q
(e.g. McElhinny and McFadden, 2000) plus its own A95 (8.1q), the suggested loop may be
anything between ca. 50q and ca. 100q. The pre-Sveconorwegian (~ 1.3 -1.0 Ga) APW
swaths of Baltica, Laurentia (including the Logan Loop) and Kalahari cratons show
similar loop-like shape suggesting that they may have drifted together (see Fig. 9 in
52
article III). However, dated paleomagnetic poles for ca. 1.25 - 1.12 Ga interval from these
continents are required to test the similarity.
In this pre-Rodinia (ca. 1.12 Ga) configuration of Baltica and Laurentia, Baltica is 90°
rotated compared to Rodinia models of Hoffman (1991), Pesonen et al. (2003) and Li et
al. (2008), but being consistent with configuration of Evans (2006a). The new VGP of
Salla dyke provides new constraint on the timing of the rotation of Baltica relative to
Laurentia (e.g. Gower et al., 1990).
Figure 18. New virtual geomagnetic poles derived from Salla diabase (1122 Ma, dotted
line) and Jänisjärvi impact melts (682 Ma, dotted line) is plotted with poles of
Baltica listed in Table 6. Shown are two different APWP curves for Baltica: (1)
after Elming et al. (1993) (solid grey line), (2) a revised highly tentative APWP
based on our new poles (dashed black line). Numbers denote the age of the pole
in Ma. R (N) reversed (normal) polarity pole.
Using the paleomagentic pole derived from the ca. 1460 Ma units of Baltica (Table 5)
and the VGP from the Salla dyke (Table 4) the proposed long-lived Mesoproterozoic
connection between Baltica and Laurentia (e.g. Åhäll and Connelly, 1998; Karlstrom et
al., 2001; Pesonen et al., 2003a and references therein) was tested (Fig. 19). We propose a
new configuration of Baltica and Laurentia in Hudsonland for ca. 1.46 Ga. This assembly
53
differs for example from those of Meert (2002) and Pesonen et al. (2003a), where eastern
Baltica has been located against eastern Laurentia. The new configuration is
reconstructed by by rotating Baltica to Laurentia’s reference frame using the Euler
parameters: Lat = 47.5°, Long = 1.5°, angle = +49° (Evans and Pisarevsky, 2008). This
particular configuration was carefully chosen for a tight cratonic fit, concordance with
basement geology, as well as pole matching (e.g. Evans and Pisarevsky, 2008). It is also
supported by geological data (e.g. Gower et al. 1990). Further evidence for this
reconstruction is the continuous rapakivi magmatism (1.65-1.42 Ga) in Baltica and
Laurentia (Rämö et al., 2005; Rämö and Haapala, 2005). The pole pairs of Baltica and
Laurentia at ca. 1.76 Ga and 1.12 Ga (see Table 4 in article III) also support this
configuration (Fig. 19).
Figure 19. Paleoreconstruction of Mesoproterozoic connection between Baltica and
Laurentia. Numbers are ages of used poles in million years. B (L) stands for
Baltica (Laurentia). Baltica and poles of Baltica have been rotated to
Laurentia’s reference frame using the Euler parameters: Lat 47.5°, Long 1.5°,
angle: +49° (Evans and Pisarevsky, 2008) Also shown rotated 1.76 Ga, 1.46
Ga, 1.27 Ga and 1.12 Ga poles for both continents by using this continental
configuration. Used data is listed in Table 4 in Article III axcept the 1460 Ma
pole for Baltica in Table 5.
4.1.2 Paleomagnetic results of the Jänisjärvi impact structure
The impact melt samples of the Jänisjärvi structure provided a stable characteristic
remanence component (Table 4). As discussed in section 3.1 samples from and near the
Jänisjärvi structure pass a positive impact test (Pesonen, 2001). Moreover, there is further
54
evidence of a Svecofennian component (component A) in various dykes on Russian
Karelia (e.g. Mertanen et al., 1999; Mertanen et al., 2006) as well as in the eastern
Finland not far from Jänisjärvi (Pesonen et al., 1991). The positive impact test (Pesonen,
2001) supports a primary nature of the Jänisjärvi ChRM. Earlier isotopic dating of the
Jänisjärvi impact melt samples yielded ages of 718 ± 5 Ma (K-Ar; Masaitis et al., 1976)
and 698 ± 22 Ma (40Ar–39Ar; Müller et al., 1990). The pole derived from the Jänisjärvi
impact melt samples (Table 4) is offset from the ~750-600 Ma APWP segment of Baltica.
At the time of writing article I, we suggested that one possible reason for such a
discrepancy could be the inherent problems in the Ar-Ar dating technique as emphasized
by Deutsch and Schärer (1994). However, Jourdan et al. (2008) provided a new highprecision 40Ar-39Ar age of 682±4 Ma (2) for the Jänisjärvi impact melt.
Figure 20. New pole derived from Jänisjärvi impactites (682, dotted line) is plotted with
poles of Baltica between ~750 and 500 Ma. Shown are two different APWP
curves for Baltica: (1) after Torsvik et al. (1996, 2001) (dotted black line), (2) a
revised APWP based on our new pole (solid grey line). Numbers denote the age
of the pole in Ma and lettering is the key for poles and references in Table 6 in
article I.
In light of these new dating results the significant deviation of the Jänisjärvi pole from the
other Neoproterozoic poles for Baltica rests on problems related to the paleomagnetic
data. Modelling of the cooling time of the melt sheet takes roughly few thousands of
years (Sazonova, 1983). This suggest that the paleomagnetic data did not sample a time
55
interval covering the longer periodicities of geomagnetic secular variation (SV).
However, the angular between-site dispersion of the poles (S = 18.2°; Table 4; Figure 4)
is consistent with the model G by Merrill and McElhinny (1983). This supports the idea
that the geomagnetic SV has been averaged. However, as already emphasized above in
the cases of the Valaam sill and the Salla dyke, matching of the S-parameter with the
model G is not conclusive evidence that SV has been averaged out. Especially, in the case
of a single intrusion (or small impact melt sheet) it is possible that the pole is a virtual
geomagnetic pole. Another possible explanation for the discrepancy rests on the fact that
the APWP for Baltica lacks well dated and well defined poles from 750 to 600 Ma as
discussed in article I. Thus, the new pole (or VGP) helps us to provide an anchor point
(682 Ma) for the APWP of Baltica. Based on our new pole, a tentative revised APWP for
Baltica is proposed (Figs. 18 and 20).
The new VGP from the Jänisjärvi impact structure places Baltica on a high latitude of ca.
60° at 682 Ma. It provides relevant data for additional tests of Laurentia-Baltica in
Rodinia. However, there is a lack of a suitable 682 Ma pole from Laurentia to_verify the
break-up of Rodinia and to reconstruct the relative positions of Laurentia and Baltica.
Table 6. Selected 1268-680 Ma paleopoles for Baltica.
Age
(Ma)
Plat
Plon
dp
dm
Vaasa dolerite
Märket dolerite
1268±13
1265±6
7
-6
164
146
3.2
7
5.5
11
Satakunta dolerite
1264±12
5
1122±7
1066±34
1042±50
11001040
995±65
158
113
293a
236
212
3
Salla diabase
Kautokeino mean
Laanila-Ristijarvi
Bamble intrusion
mean
Årby dolerite
2
71
-71a
-36
-2
8
3
13
8
5
21
3
-7
217
227
15
7
15
10
Nilstorp dolerite
966±2
9
239
5.8
10.6
Dalarna (4 dykes)
Göteborg-Slussen (4 dykes)
Hunnedalen (4 dykes)
Neoproterozoic meanb
946±1
935±3
848±27
750
5
-7
-41
28
45
-45a
239
242
222
197
77
257a
15
12
11
8
15
12
12
8
this work
Mertanen et al., 1994
Mertanen et al., 1996
Brown and McEnroe, 2004; Torsvik and
Eide, 1998
Patchett and Bylund, 1977
Patchett and Bylund, 1977; Söderlund et
al., 2004
Bylund, 1992;
Söderlund et al., 2005
Pisarevsky and Bylund, 2006
Walderhaug et al., 1999
Meert and Torsvik, 2003
10
10
this work
Rock unit
Jänisjärvi impact melt
a
682±4
Reference
Neuvonen, 1966;
Neuvonen and Grundström, 1969
Neuvonen, 1966;
b
Reversed polarity pole, Pole polarity inverted
56
4.2 Vredefort impact structure
In article IV, we present new paleomagnetic and rock magnetic results from the Vredefort
impact structure and Johannesburg dome (South Africa). First, we will discuss
paleomagnetic results and their implications on Paleoproterozoic reconstructions. Finally,
the observed high Koenigsberger ratios (Q values) will be discussed.
4.2.1 Paleomagnetic results of the Vredefort impact structure – implication for the
paleoreconstruction of the Kaapvaal Craton at 2.0 Ga
The Vredefort impactites provided a stable hard coercive and high-temperature ChRM
component (D = 18.3°, I = 54.8°, 95 = 8.1°). This result is consistent with earlier
paleomagnetic studies by Hargraves (1970), Hart et al. (1995) and Carporzen et al.
(2006). A similar component, carried by magnetite, with the direction of D = 2.1° I =
62.5° (95 = 7.2°) is also isolated in some clasts of Archean basement complex samples.
Hart et al. (1995) also separated a similar impact-related component (D = 25°, I = 56°, 95
= 16°) from the basement samples. The fact that an impactite-type direction is separated
also from the Archean rocks and that Archean direction can be observed in samples from
both Vredefort and Johannesburg domes, indicates that the magnetization of impactites is
of primary origin (a positive impact test; Pesonen, 2001). Added to this, a similar
component has been seen in various lithologies from the Kaapvaal Craton (e.g. Morgan
and Briden, 1981; Jackson, 1982; Layer, 1986; Hart et al., 1995; Maré et al., 2006) and
are interpreted to be related to two known ca. 2.0 Ga events (the Vredefort impact event
and the Bushveld Igneous Complex). The observed ChRM provides a well dated (2.023
Ga) pole of 25.4°N, 43.8°E with A95 = 9.6°, which places the Kaapvaal Craton on the
paleolatitude of 35°. The mean pole for Vredefort impactites falls on the Paleoproterozoic
part of the APWP of the Kaapvaal craton.
We hoped that paleoreconstruction of the continent configuration at 2.0 Ga would shed
some light on the time interval between the proposed supercontinents Kenorland and
Hudsonland. However, there is not enough paleomagnetic data to predict whether longlived near-equatorial Kenorland (Evans et al., 2002) was a true supercontinent (Pesonen
et al., 2003a). It has been proposed that this oldest supercontinent broke-up during
protracted (or episodic) rifting ca. 2.45–2.15 Ga ago (e.g. Heaman, 1997; Aspler and
Chiarenzelli, 1998; Bekker et al., 2001). It is known that Fennoscandia and Superior
experienced another rifting at 2.20-2.00 Ga, which is evidenced by sedimentation and
widespread mafic dyke activity (e.g. Vuollo et al., 1995; Melezhik et al., 2007). Here we
propose a model for the Kenorland configuration based on data listed in Table 7 (Fig. 21
a). Unfortunately, there is no quality published constraint on the position of Kaapvaal at
2.45 Ga. In this model, Superior lay at near-equatorial latitudes, whereas Australia
(Yilgarn) and Fennoscandia were at much higher (> 30° southern) latitudes. Karelian
(Mertanen et al., 1999) and Matachewan (Bates and Halls, 1990) dyke swarms have
similar trends in this proposed configuration. Similar lithostratigraphic sequences have
been documented in Superior and Fennoscandia during 2.45 and 2.20 Ga (Bekker et al.,
2001), which support their proximity. Between 2.40 and 2.22 Ga, Superior and
57
Fennoscandia experienced one to three successive glaciations (e.g. Bekker et al., 2001)
and the glaciogenic sequences also contain paleoweathering zones (regoliths or paleosols)
lying generally on top of the glaciated layers. Evans (2003) notes that these are
lithostratigraphically similar to the Neoproterozoic strata that also contain glaciogenic
layers and paleoweathering zones. In both case (Kenorland and Hudsonland), the
paleomagnetic data point to low and even equatorial latitudes (<45°) during glaciations
(e.g. Evans, 2000; Evans et al., 1997; Williams and Schmidt, 1997).
Figure 21. a) The reconstruction of continents at 2.45 Ga. Data available from Superior,
Fennoscandia and Australia (Table 7). Karelia (Baltica), Superior (Laurentia)
and West Australia (Australia) are shown in dark shading. Dyke swarms are
shown as sticks and they are: Matachewan (Laurentia), Russian Karelian
(Baltica) and Widgiemooltha (West Australia), respectively. b)The
reconstruction of continents at 2.0 Ga. Data available from Kaapvaal, Superior
(Laurentia), Fennoscandia (Baltica), Amazonia, West Africa, Ukraine, and
Congo/Saõ-Francisco (Table 7). The shaded area in the West Africa Craton
represents mainly Eburnean cratonized areas (the Man block to the right and
the Reguibat block to the left) and in the Amazonia Craton they represent the
Guyana Shield (to the right) and the Guaporé Shield (to the left), respectively.
To obtain the position of Kaapvaal in the 2.0 Ga configuration we used a combined pole
of Vredefort, Lower Swaershoek Formation, and Phalaborwa 1 data (Table 7; Fig. 21b).
For other continents except Baltica (e.g. Fennoscandia) we used poles listed in Pesonen et
al. (2003a). The position of Fennoscandia is based on the 2058 Ma data from Kuetsyarvi
lavas (Torsvik and Meert 1995; Evans and Pisarevsky, 2008). We want to emphasize that
such a reconstruction is only tentative due to the fact that the ages of used paleomagnetic
poles are not coeval among different cratons (Table 7). However, the configuration
shown in Fig. 21 emphasizes the independent drift of continents at ca. 2.0 Ga rather than
58
pointing to a large landmass as suggested by Pesonen et al. (2003a). It is noteworthy that
continents occupy low to intermediate latitudes at 2.0 Ga as they did during both
Kenorland and Hudsonland assemblies (e.g. Evans et al., 2002; Pesonen et al., 2003a).
Table 7. Paleomagnetic poles used for reconstruction.
Craton
Formation
Age (Ma)
Plat
(°N)
Plon
(°E)
A95
(°)
References
Reconstruction at 2.45 Ga
B
Combined Burakovka intrusion
and associated dykes and Lake
Pääjärvi D-comp.
2449 ± 1
-16.9
251.6
-
Mertanen et al., 2006
L
Matachewan dykes
2446-2473
-42.0
238.0
3.0
Bates and Halls, 1990
Au (WAu)
Widgiemooltha
2418-2420
-9.0
157.0
8.0
Evans, 1968;
Nemchin and Pidgeon, 1998
2023 ± 4
25.1
43.5
9.6
Reconstruction at 2.0 Ga
K
Vredefort
K
Vredefort
K
K
K
K
K
Vredefort
Vredefort
Lower Swaershoek Formation
Phalaborwa 1
Mean Kaapvaal
2023 ± 4
2023 ± 4
2023 ± 4
2054 ± 4
2060.5 ± 0.6
2061-2023
21.6
21.4
22.2
36.5
35.9
27
46.6
27.1
49.1
51.3
44.8
44
6.2
15.7
20.4
10.9
8.5
12
this work
Carporzen et al. 2006,
recalculation this work
Hargraves 1970
Hart et al. 1995
de Kock et al. 2006
Morgan and Briden 1981
this work
L
Minto dykes
1998 ± 2
38.0
174.0
10.0
Buchan et al., 1998
Am
Mean Amazonia
2019
42.0
181.0
-
Pesonen et al., 2003a
C-SF
Uauá dykes
1983 ± 31
2200 ±23
24.0
331.0
4.0
D’Agrella-Filho and Pacca,
1998; Bastos Leal et al., 1994
WA
Liberia M
2050 ± 6
-18.0
89.0
13.0
Onstott and Dorbor, 1987
U
Mean Ukraine
2000
53.0
142.0
-
Pesonen et al., 2003a
B
Kuetsyarvi lavas (Karelia )
2058 ± 6
23
298
7
Evans and Pisarevsky, 2008
Plat, Plon, latitude and longitude of the paleomagnetic pole. A95, the 95% confidence
circle of the pole. K Kaapvaal, L Laurentia, B Baltica, Am Amazonia, C-SF Congo-Saõ
Francisco, WA West Africa, U Ukraine, Au Australia (WAu West Australia).
4.2.2 Younger ca. 1.1 Ga magnetization
Detailed paleomagnetic analysis of all rocks, including the Vredefort impactites, yielded
occasionally distinct great circle paths either towards a shallow north or a shallow south
direction (D = 160°, I = -5°, 95 = 21°). As this direction is also isolated from the baked
59
zone of the Anna’s Rust Sheet gabbro samples (1108.6 ± 1.2 Ma; Hanson et al., 2004b) in
the northern collar of the Vredefort structure, we interpret this direction as an overprint
caused by the 1.1 Ga magmatic events, known as the Umkondo Large Igneous Province
(Hanson et al., 2004a; 2004b). Such great circle trends were also reported in several
gabbro, wehrlite, pyroxenite and granulite sites (McDonald and Andersen, 1973), and in
the Archean basement granitoids (Jackson, 1982).
4.2.3 Archean magnetization
The Archean basement (ABC) samples from the Vredefort and Johannesburg domes with
unusually high Q–values show variable directions of hard coercivity magnetization (Fig.
22). This has also been reported in earlier paleomagnetic studies (Jackson, 1982; Layer,
1986; Layer et al., 1988, 1989a, 1989b; Hattingh, 1989; 1999; Hart et al., 1995, 2000;
Carporzen et al., 2005). However, the majority of our specimens show congruent
directions of this high coercivity magnetization within one sample, which contrasts the
observation of Carporzen et al. (2005) and Hart et al. (2000), who reported variable
directions at the centimeter scale in all of their samples. Also, samples with normal Qvalues show a slight trend towards steep magnetization directions (see Fig. 22b) which is
consistent with Layer et al. (1989a; 1989b). The mean of the direction is (D = 345°, I =
63°, 95 = 10°) and yields a paleomagnetic pole of Plat = -18°, Plon = 196° (A95 = 13.4°),
for results from 14 samples (56 specimens). A steep direction for Archean magnetization
of has also been noted by Layer (1986).
Figure 22. Mean directions of the samples from the Archean Basement Complex (ABC) of
the Vredefort (V) and Johannesburg (JHB) domes. a) Magnetization direction of
samples with high values of NRM intensity (high Q), and b) magnetization
direction of samples with “normal” values of NRM intensity (normal Q). Mean
direction for samples inside the dotted line is marked with a rectangle.
60
4.2.4 High Q values and the origin of remanent magnetization
In article IV, an important rock magnetic result is that high Q values in Archean samples,
also reported by previous studies (e.g. Jackson, 1986; Hart et al., 1995; 2000 and Pesonen
et al., 2002; Carporzen et al., 2005; 2006), are now seen for the first time in the Archean
rocks of the Johannesburg Dome. This rules out the direct Vredefort impact connection as
an explanation for high Q-values (Carporzen et al., 2005). Such an interpretation is also
supported high Q values which are independent of the distance from the center of the
structure (see Fig. 16 in article IV). This contradicts Hart et al. (1995), who proposed that
these values were dependant on the distance from the center of the structure.
We observed that several impactite samples also showed elevated NRM values (high Q),
which cannot be explained by extremely rapidly occuring plasma-induced magnetic field
caused by the impact as was suggested by Carporzen et al. (2005). The impact rocks were
cooled more slowly (several hundreds of thousands of years; e.g. Ivanov, 2005; Ivanov
and Deutsch, 1997) than the lifetime of a plasma field (from seconds to minutes; e.g.
Crawford and Schultz, 1999; Turtle and Pierazzo, 1998). Another explanation for the
high Q values could be lightning. However, we discount lighting as the cause of all
observed high Q values (expect in the case of ARS samples and some of the ABC
samples) based on the Lowrie-Fuller-test and the fact that all the studied lithologies from
all the surface exposures, as well as some samples (ABC and PT breccia) from the
quarries, show elevated Q values. Moreover, samples with high Q-values showed
relatively hard magnetization with multiple remanent magnetization components, which
contrasts lightning-induced magnetization (soft and normally single-component; e.g.
Irving, 1964). As we exclude lightning as a cause, high Q values may be related to the
circulations of hot fluids taking place within the two domes (Vredefort and
Johannesburg). In the case of Vredefort, the impact event is most likely the triggering
cause for the uplifting (doming) of the middle crustal rocks with enhanced (high Q)
magnetization. For Johannesburg, another, yet unknown, explanation for the doming
must be found.
The observed high Q-values is a puzzling feature, and some novel experiments must be
carried out to solve the cause. We suggest a few possible experiments to study the
behaviour of magnetization of Archean rocks from these domes: (1) simulate the impact
process using plastic explosive (velocity 8.2 km/s) on the rocks to study how the shock
affects the magnetization of the Archean rocks of both Vredefort and Johannesburg
dome; (2) carry out artificial lightning experiments using both high voltage and high
current experimental set-ups.
5. Conclusions
The first task of this work was to obtain new Neo- to Mesoproterozoic paleomagnetic
data for the apparent polar wander path of Baltica and to study proposed long-lived (1.82
– 1.27 Ga) connection between Baltica and Laurentia. We were able to obtain three new
well dated poles for Baltica: (1) a 1458 +4/-3 Ma virtual geomagnetic pole (VGP) of Plat
61
= 14°N,and Plon = 166°E, (A95 = 2.4°) from the Valaam sill, combined with other ca.
1460 Ma results from Baltica to produce a robust paleomagnetic pole of Plat = 11°N,
Plon = 173°E, (A95 = 7.5°) was obtained; (2) a 1122±7 Ma VGP of Plat = 71°N,and
Plon = 113°E (A95 = 8.1°) from the Salla dyke: (3) a 682±4 Ma VGP of Plat = 45°N, and
Plon = 77°E (A95 = 10.4°) from the Jänisjärvi impact structure. Based on poles from Salla
diabase dyke and from Jänisjärvi impactites tentative revised segments of apparent polar
wander paths for Baltica are proposed.
A new configuration of Baltica and Laurentia within the supercontinent Hudsonland was
proposed based on the observed 1460 Ma paleomagnetic pole and the VGP of the Salla
dyke. Furthermore, these provide evidence for a long-lived (1.76 – 1.12 Ga) Baltica Laurentia connection. The new VGP of Salla dyke indicates that Baltica rotated relative
to Laurentia already at 1.12Ga.
The second objective of the work was to study the suitability of magnetically stable
samples from two impact structures (the Jänisjärvi structure (682 Ma) from Russia and
the Vredefort structure (2.023 Ga) from South Africa) for paleomagnetic purpose.
Paleomagnetic and rock magnetic studies of these samples show that impact rocks carry a
stable magnetization, presumably of thermal origin. The primary nature of magnetization
in impact generated rocks was verified using the impact test (Pesonen, 2001). Such data
provide a useful tool to fill the gaps in Precambrian apparent polar wander paths and thus
helps us to obtain a more continuous record of the positions of continents during the past.
The third objective was to provide new poles for the Paleoproterozoic apparent polar
wander path of the Kaavaal Craton using samples from the Vredefort impact structure
and it’s surroundings. The Vredefort impactite samples pass the impact test and provide a
well dated (2023 ± 4 Ma) pole of Plat = 25.4°N and Plon = 43.8°E (A95 = 9.6°) for the
APWP of the Kaapvaal craton. Such a pole is consistent with the previous studies.
The fourth objective of this work was to study the unusually high Koenigsberger ratios
(Q values) observed in various lithologies in and around the Vredefort structure. This was
done using numerous rock magnetic experiments coupled with paleomagnetic data.
High Q values, which had also been reported by previous studies, are now also seen in
samples from the Johannesburg Dome, where there is no evidence for impact, thus a
direct causative link to the Vredefort impact event can be ruled out. As the observed
magnetization is hard and shows multiple components of remanent magnetization, we
exclude lightning as a cause for the high Q values (except in case of a 1.1 Ga gabbro).
The high Q values may instead be related to the circulation of hot fluids taking place
within the two domes (Vredefort and Johannesburg). In the case of Vredefort, the impact
event is most likely the triggering cause for the uplifting (doming) of the middle crustal
rocks with enhanced (high Q) magnetization. For the Johannesburg dome another, yet
unknown, explanation for the doming must be found.
62
6. References
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V.A.V., 1994. Geocronologia Rb/Sr e K/Ar do enxame de diques ma´ficos de Uaua´,
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