Paleomagnetism and Applications to Structural Geology
Aubrey Weese
Environmental Science and Public Policy,
George Mason University, Fairfax VA
April 28, 2003
Paleomagnetism is the study of the direction and intensity of the earth's magnetic field through geologic time. It is useful for determining the movement of rocks both on a large and small scale. On a large scale, paleomagnetism helps show how the continental and oceanic plates have drifted relative to the earth's spin axis and to one another. On a small scale, it helps to determine the movement of crustal blocks in continents, particularly vertical axis rotation. Finally, it can also be used to temporally correlate rocks on any scale (a study called magnetostratigraphy).
In this paper, I will discuss the general properties of the earth's magnetic field, how this magnetism is imprinted upon rock bodies, and finally how this imprint is deciphered by scientists to unravel their history.
The Earth's Magnetic Field
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The primary component (more than 90%) of the magnetic field observed
at the earth's surface is generated internal to the planet by electric currents
flowing in its liquid outer core. This field varies slowly with time and
space, and can be described by mathematical models such as the International
Geomagnetic Reference Field and the World Magnetic Model.
At any point on the surface, the earth's magnetic field can be represented by a vector with a specific length and direction. When averaged over a long period of time (greater than 105 years), this field can be represented by a dipole magnet at the earth's center with its north-south axis being the same as the spin axis of the earth. |
Therefore, according to this model, a rock that acquired its magnetization at the equator would have a time-averaged vector inclined at 0° to horizontal. Similarly, a rock that acquired its magnetization at the north pole would have a vector inclined 90° to horizontal. Those in between have a vector described by the equation tan I = 2 tan l, where I is the inclination and l is the latitude. So, if rocks on a continent that is currently situated 20°S have a magnetic vector that reads 0°, it shows that the continent must have moved 20° to the south since the formation of those rocks. East-West motion (about lines of longitude) is undetectable using this method (Access Science 2001).
Magnetism in Rocks
Magnetization in rocks is due to the magnetic moments of neighboring atoms in minerals in the rock being coupled parallel or antiparallel. These magnetic minerals usually make up a few percent or less of the rock. One of the most common ones is magnetite Fe3O4. The figure below shows its crystal structure.
(Moskowitz 1991) |
A site: Fe ion is surrounded by four oxygensB site: Fe ion is surrounded
by six oxygensThe magnetism is due to the spin on the electrons in the A
sites being antiparrell to the spin on the electrons in the B sites.
Another common magnetic mineral is hematite Fe2O3. In the presence of no magnetic field, these minerals have a spontaneous magnetization. This magnetization consists of different regions of the crystal having their magnetic vector going in a different random direction. These regions are called domains. The presence of a large magnetic field can enlarge or reorient these domains in such a way that when the field is removed, the original domain sizes and orientations are not recovered. The rock now has a permanent magnetization, known as remnant magnetization (Access Science 2001). The magnetic field of the earth is not strong enough to produce remnant magnetization except in special circumstances, which vary according to rock type. Igneous rocks become susceptible to magnetization by the earth's field once they are heated to a special temperature known at the curie temperature. This temperature is different for different minerals (for example magnetite = 580°C and hematite = 670°C). |
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As the rocks cool to a slightly lower temperature, known as the blocking
temperature, the magnetization becomes permanently locked within them, and
will remain unless the they are reheated to near the blocking temperature.
Blocking temperatures for magnetite-bearing rocks range from 500-580°C.
Magnetization acquired by this kind of heating and cooling is known as
thermoremnant magnetization (TRM). It is often acquired in several steps
throughout the slow cooling history of a rock, as illustrated in the figure
on the left.
Sedimentary rocks acquire their magnetization as magnetic grains fall through the water and accumulate with other sediments. These grains can become aligned with the ambient magnetic field as they are deposited, resulting in depositional remnant magnetization (DRM). In addition, chemical changes in the rock during the process of lithification can cause new magnetic minerals to grow. Once these minerals reach a critical grain size (about 0.2 micrometers for hematite) they will acquire a chemical remnant magnetization (CRM). These original magnetizations can be partially overprinted by subsequent geologic events, such as metamorphism, hydrothermal alteration, small influences of the earth's more recent magnetic field acquired over long periods of time (viscous remnant magnetization), and electrical currents caused by lighting strikes (isothermal remnant magnetization) (Access Science 2001). |
Paleomagnetic Methods
The goal of paleomagnetic studies is to separate and successively erase the different layers of remnant magnetization in rock units, associating an age with each. The first step is to collect carefully oriented samples. These samples are usually drilled out of rocks using portable diamond coring equipment. Then the samples are taken to the lab where they are cut into small upright cylinders, and their magnetization is measured using a sensitive magnetometer. In the past astatic or high speed spinner magnetometers were used; nowadays either flux-gate detectors or cryogenic detectors are commonly used (Access Science 2001).
The next step is to subject the samples to a series of demagnetization steps of increasing severity, remeasuring the magnetization after each step. During this processes the magnetization vector will change until the most stable component has been isolated, at which point the vector will decay in a straight line to the origin. This final component is called the characteristic remnant magnetization (ChRM) (Butler 1992).
Several types of demagnetization techniques can be used in this process, each having a slightly different purpose. In alternating field demagnetization, the sample is put into an enclosure with zero magnetic field and then exposed to an alternating sinusoidal field with smoothly decreasing amplitude. This procedure progressively randomizes the magnetic grains whose coercive force is equal to or less than the peak amplitude of the alternating field. This technique is useful for removing IRM and for resolving one generation of TRM from another. In thermal demagnetization, rock samples are heated to a specific temperature and then cooled in an enclosure with zero magnetic field. By repeating this process at progressively higher temperatures, magnetizations with different blocking temperatures can be removed. This technique is also useful for separating TRM episodes from one another as well as for removing recently acquired VRM. Finally, chemical demagnetization is done by immersing rocks in concentrated hydrochloric acid for periods of time lasting from days to weeks. This acid dissolves the precipitated magnetic minerals in the rocks from the outside inward (which is the reverse order of how they grew). This technique is useful for separating different CRM episodes in all rocks, and both DRM and CRM episodes in sedimentary rocks (Access Science 2001).
Once a paleomagnetic study has been completed, the researcher will have a list of vector directions from the collected samples. These directions are composed of the magnitude B, the inclination I (the angle that the vector makes with the horizontal) and the declination D (the angle that the vector makes relative to true north, measured positive clockwise from north). Each vector is then rotated from sample coordinates to geographic coordinates, and possibly corrected for the tilt of the bed that the sample is in.
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If the samples gained their magnetism after the bed was titled, they
do not need to be corrected this way, but if they were magnetized when the
bed was in its original horizontal state, then they do. The paleomagnetist
can determine which is the case by applying Graham's fold test. This test
is done by plotting both the original data and the tilt-corrected data on
a stereonet. If the data points cluster closer together after they are
tilt-corrected, then the magnetization was most likely acquired pre-tilting.
If on the other hand, the tilt correction causes them to be more spread apart,
the magnetization was likely acquired post-tilting. See the figure on the
left for an example of both cases (Tauxe 2002).
If the data points are most tightly clustered somewhere between these two states (say, when they have been only 50% tilt-corrected), this can be an indication of magnetization acquired syn-folding (sometime during the folding process). In fact, sedimentary rocks such as shallow water carbonates often only carry magnetizations acquired during tectonic folding events. However, partial tilt-correction clustering on stereonets can also be due to other factors, such as vertical-axis rotation. So a careful look at the structure and history of the region is required to determine this for sure (Weil et al. 2002). |
In addition to the fold test, several other tests can be applied to get an idea for when the rocks were magnetized. The conglomerate test can be used if the rock type that is being studied is found in a conglomerate bed. If the magnetic vectors of that rock type in the bed have an ordered and tightly clustered orientation, yet the other clasts in the bed have a randomly dispersed orientation, then the rocks being studied must have been magnetized prior to the formation of the conglomerate. The baked contact test can be used if the rock type that is being studied is found in an area with igneous intrusions. The intrusion will heat the surrounding rock to the curie temperature, therefore imparting its magnetization to the adjacent rock in a zone that decreases with distance from the intrusion. So if rocks right next to the intrusion have different magnetic vectors than those far way, we know that magnetization happened before the intrusion. But if the rocks everywhere have the same magnetic vector than it must have happened afterwards (Tauxe 2002).
Apparent Polar Wander Paths
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Over time periods shorter than 105 years, the earth's magnetic
field shows large local variations. These are termed secular variations.
If the drilled samples cover a long enough time period, their vectors can
be averaged together to eliminate secular variations and get a representative
paleomagnetic pole for the rock unit. Roughly one hundred sites (each with
nine to ten samples), spanning 100,000 years are needed to sufficiently average
out secular variations (Access Science 2001). The direction of magnetic vectors
measured from these sites should have a dispersion ranging from around 10
to 25 degrees. The chart on the left shows typical dispersions due to secular
variation at different latitudes. The dispersion increases by almost a factor
of two from equator to pole.
If the data points are clustered tighter than this, most likely a longer time interval needs to be sampled. If on the other hand they are scattered wider than this, it indicates an additional source of dispersion besides just secular variation. This dispersion could be due to tectonic disturbance or difficulty calculating the ChRM's. Both cases cast doubt on the reliability of the data (Butler 1992). |
Paleomagnetists use fisher statistics in order to calculate a mean vector from all these samples. This method was developed for studying the dispersion of unit vectors on a sphere. The figure below shows a scattered (a) and clustered (b) dataset, with their fisher mean vector indicated by a star and the 95% confidence bounds indicated by a circle around this star (Tauxe 2002).
(Tauxe 2002)
There are a number of factors that lead to scatter in paleomagnetic data.
uncertainty in the measurement caused by instrument noise or sample alignment errors,
uncertainties in sample orientation,
uncertainty in the orientation of the sampled rock unit,
variations among samples in the degree of removal of a secondary component,
uncertainty caused by the process of magnetization, and
secular variation of the Earth's magnetic field.
(Tauxe 2002)
Items number 1, 2 and 6 above lead to a symmetric distribution about a mean direction, which is a fischer distribution. For the most part, paleomagnetists assume this kind of distribution for their data. However in some instances this is not a correct assumption and different methods need to be used to calculate the mean vector. For example, number 4 on the list above can lead to an asymmetric distribution of data points that are smeared about a great circle. Also, if the data comprises both normal and reversed polarity states, it can have a bimodal distribution (Tauxe 2002). According to Borradaile et al. (2003) there are also several factors that can lead to an elliptical distribution (termed a bingham distribution) instead of a spherical distribution. These factors are caused by metamorphism, and include: magnetic anisotropies in the rock caused by foliation, deflection of remanence vectors by finite strain towards the maximum extension direction, and stress induced remagnetization by mobilization of domain walls.
In these cases, an alternative method statistical method is needed, such as bingham statistics or a statistical bootstrap. The figure below shows a dataset that would require alternative methods along with bootstrap calculated mean vectors and confidence circles (Tauxe 2002).
(Tauxe 2002)
Once the averages are calculated, poles for units of different ages can then be put into an ordered time sequence and connected on a map, creating an apparent polar wander path (APWP). This path shows how the axis of the earth has appeared to move over time relative to the rocks being sampled (though it is actually the continent that the rocks are on that is doing the moving). Therefore these paths can be interpreted to show the displacement history of continents and tectonic plates. In fact they are the only direct method we have to obtain relative motion history for the continents at times earlier than about 2.5 x 108 years before present (Access Science 2001). In the figure below, you can see that the paths form an arc tracking away from the geographic north pole in the northern hemisphere. Because of magnetic reversals, they form a similar arc tracking away from the geographic south pole in the southern hemisphere (Butler 1992).
(Access Science 2001)
APWP are well known for some geologic time intervals and poorly known for others. As more data is collected and research methods become more accurate, they are continuing to be improved. As you can see from the picture, the paths are separated by some sharp turns or cusps. Some researchers consider the straight tracks as periods of time when the lithospheric plate was rotating about a fixed point called a euler pole. The cusps indicate times of reorganization of plate boundaries and tectonic forces.
The matching of APWP is the primary paleomagnetic method used for proposing and testing past relative positions of continents. If they were all rotating about the same euler pole, then their APWP's should coincide. The picture on the left above shows the Mesozoic and Paleozoic paths for north America (solid circles) and Europe (open circles). In the picture on the right the continents have been rotated until the two paths coincide. Theoretically, this map then maps their position back in the Mesozoic and Paleozoic (Butler 1992).
Application to Regional Tectonics
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The second aspect of paleomagnetism I would like to discuss is its
application to regional tectonics. This is perhaps of most interest to structural
geology. In addition to describing the motion of continents, paleomagnetism
can also be used to determine the motion of pieces of the continent known
as terranes. A terrane is a fault bounded region that has a geologic history
distinct from neighboring regions. Dimensions of terranes are up to hundreds
of kilometers (Butler 1992).
Motions of these crustal blocks can be calculated by comparing their observed magnetic inclination and declination to the expected declination and inclination they would have at that latitude. If the inclination angle is flatter than expected, it indicates latitudinal motions towards the paleographic pole. Changes in the expected declination angle indicate vertical axis rotation of the crustal block. Both of these changes occur together as the block rotates about a euler pole in the same way that continents rotate about one. The inclination flattening, F, is given by F = Ix - Io and the rotation of declination, R, is given by R = Do - Dx. (R is defined as positive when Do is clockwise of Dx). This method is known as the direction-space approach and is illustrated in figure (a) to the left. (Butler 1992) Another method is pole-space approach, where the apparent polar wander path of the continent serves as a reference direction to determine how the crustal block has moved relative to the entire continent. In this approach, comparison is made between the reference pole of the continent (RP) and the observed pole of the crustal block (OP) located at position S. This method is illustrated in figure (b) to the left. A spherical triangle is drawn with coordinates at S, OP and RP. The vertical axis rotation, R, is the angle of the spherical triangle at S. If the angular distance from S to OP is po and the angular distance from S to RP is pr, the poleward transport is given by p = po - pr. p is positive if the block has moved towards the reference pole and negative if it has moved away from the pole (Butler 1992). |
A few cases where this kind of study has been done are as follows:
Jones et al (1976) noted that structures in the Mesozoic rocks of Transverse Ranges of southern California are oriented east-west, yet similar structures in the Mesozoic rocks from Oregon to Baja California are oriented north-south. This led them to suspect a vertical-axis rotation. This was confirmed by Kamerling and Luyendyk in 1979, who used paleomagnetic data collected from the Conejo volcanics exposed in the Santa Monica mountains and Conejo hills to show a that a major clockwise rotation had occurred in the western Transverse Ranges.
In 1957 Cox observed that the paleomagnetic declination in the Eocene Siletz River volcanics of the Oregon Coast range was east of what would be expected. In 1977 he confirmed together with Simpson that the Oregon Coast range had rotated clockwise about 70° since the Eocene. In 1988 Wells and Heller extended this study to determine that this rotation was due to a combination of right shear between the oceanic plates and the North American plate (40%) and northwards decreasing amount of extension in the Basin and Range Province east of the Cascade Arc (60%).
In a more recent study, Pueyo et al. (2003) used paleomagnetic data to demonstrate a vertical axis clockwise rotation of around 32° in the western External Sierras in the southern Pyrenees mountain range of France. This rotation is due to conical syncline that developed as a result of flexure of the foot wall ramp in the south Pyrenean sole thrust fault.
From these examples, it is clear that paleomagnetism is an effective way to determine vertical axis rotation. This process is very difficult to detect by other methods, yet a growing list of examples is showing that it is a major tectonic process in continental deformation.
Magnetostratigraphy
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Another property of the earth's magnetic field is that it periodically
changes in polarity. The axis of the dipole remains the same, but the north
end switches to the south end and the south end switches to the north end.
On average, the earth spends about half its time in each state. Typically,
the transition between states takes around 1000-2000 years (instantaneous
relative to geologic time). The time span between these switches, however,
can very greatly. This means that the sequence of normal (N) and reverse
(R) polarity bands in rocks can be highly distinguishing when doing correlations
(Access Science 2001).The first step when doing this kind of correlation
is to consult a geomagnetic polarity time scale. For periods over the past
5 million years, K-Ar dating and paleomagnetic studies are used to develop
these scales. Earlier than that, the uncertainties in K-Ar dating are as
large as the duration of magnetic polarity events, so alternative methods
have to be used, the primary one being paleomagnetism of deep-sea cores (Butler
1992). For the period of about the mid-Mesozoic to the present, the richest
source of polarity time scale information comes from basaltic oceanic crust
formed at mid-oceanic ridge spreading centers. This crust cannot be directly
sampled, but it is possible to remotely sense its polarity. Normal polarity
rocks on the sea floor produce a field that adds to the regional geomagnetic
field, producing a positive magnetic anomaly. Reversed polarity rocks subtract
from the regional field, producing a negative anomaly.
Therefore, a boat towing a magnetometer that measures the magnetic field at the sea surface can be used to create a magnetic anomaly profile that will correspond with the reversals recorded in the rocks below. Rocks become progressively older as they extend outwards from a mid oceanic ridge. By comparing their reversal records with those measured by other means, it is possible to determine the rate of sea floor spreading. From this data, the approximate age of the sea floor crust can be calculated, which lead to the development of the magnetic reversal chart shown on the previous page (Butler 1992). By mapping out the polarity zones in a sedimentary stratigraphic section and comparing them with this scale, the age of events in the section can be determined. In these kind of studies, unambiguous determination of the polarity of the ChRM data for the rocks is the most important goal of the paleomagnetist. Fine grained rocks (mudstones, sandstones, siltstones) are preferred for this because they acquire DRM more efficiently than coarse grained rocks. They are also less susceptible to the acquisition of secondary CRM because they are less permeable. A minimum of three to four samples per site should be collected, and the sites should be spaced close enough together to be sure that reversals in the section are not skipped. The entire section should cover a time span of at least 2 million years (Butler 1992). A list of the criteria needed for a reliable magnetostratigraphic correlation is given below, followed by an example image of what such a correlation might look like.
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(Tauxe 1998)
Conclusion
From the varied applications that have been discussed, one can see that paleomagnetism is indispensable to many fields of geology. Apparent polar wonder paths and magnetostratigraphy of the sea floor were the initial evidence that proved the theory of continental drift and sea floor spreading. They continue today to be one of the best, and in same cases only way to accurately determine the past location of continents. In addition, regional paleomagnetism is the only accurate way we have of determining vertical axis rotation, which is a key component in deformation.
There are paleomagnetism labs in universities all over the world and frequent articles on it are published all the time in physics and geology journals, discussing new research findings and improvement of methods. The field is definitely vital and growing.
References
Borradaile, Graham J., Tomasz Werner, and France Lagroix, Differences in paleomagnetic interpretations due to the choice of statistical, demagnetization and correction techniques: Kapuskasing Structural Zone, northern Ontario, Canada, Tectnophysics, v. 363, 103-125, 2003.
Butler, Robert F. PALEOMAGNETISM: Magnetic Domains to Geologic Terranes. Blackwell Scientific Publications, 1992.
Cox , A. V., Remanent magnetization of lower middle Eocene basalt flows from Oregon, Nature, v. 179, 685-686, 1957.
Jones, D. L., M. C. Blake, and C. Rangin, The four Jurassic belts of northern California and their significance to the geology of the Southern California borderland, In: Aspects of the Geologic History of the California Continental Borderland, ed. D. G. Howell, Misc. Publ. 24, Am. Assoc. Petrol.. Geol., Tulsa, Okla., pp. 343-362, 1976.
Kamerling, M. J. and B. P. Luyendyk, Tectonic rotations of the Santa Monica Mountains region, western Transverse Ranges, California, suggested by paleomagnetic vectors, Geol. Soc. Am. Bull., v. 90, 331-337, 1979.
Moskowitz, Bruce M. "Hitchhiker's Guide to Magnetism" Institute for Rock Magnetism. University of Minnesota. June 1991. http://www.geo.umn.edu/orgs/irm/hg2m/hg2m_index.html
"Paleomagnetism" and "Rock Magnetism" Access Science Encyclopedia. McGraw-Hill Companies. 2001. http://www.accessscience.com/
Pueyo, E.L., J.M. Pares, H. Millan, A. Pocovi, Conical folds and apparent rotations in paleomagnetism (a case study in the Southern Pyrenees), Tectonophysics, v. 362, 345-366, 2003.
Simpson, R. W. and A. Cox, Paleomagnetic evidence for tectonic rotation of the Oregon Coast Range, Geology, v. 5, 585-589, 1977.
Tauxe, Lisa. "Paleomagnetism." Prepared GSA Short Course. 1998. http://sorcerer.ucsd.edu/es160/lecture8/web6/
Tauxe, Lisa. "SIO 247: Rocks and Paleomagnetism". Lecture Notes. Scripps Institution of Oceanography. 2002. http://magician.ucsd.edu/~ltauxe/sio247/
Weil, Arlo B. and Rob Van der Voo, The evolution of the paleomagnetic fold test as applied to complex geologic situations, illustrated by a case study from northern Spain, Physics and Chemistry of the Earth, v. 27, 1223-1235, 2002.