Chapter 1 - Plate Tectonics

show that simultane-
ous crystallization of some Fe-Ti oxides with different 
Curie temperatures can cause these minerals to become 
magnetized with an opposite polarity to the ambient field. 
This self-reversal magnetization is related to ordering 
and disording of Fe and Ti atoms in the crystal lattice. 
Although self-reversal has occurred in some young lava 
flows, it does not appear to be a major cause of reverse 
magnetization in rocks. The strongest evidence for this 
comes from correlation of reverse magnetization between 
different rock types from widely separated localities. For 
instance, reversed terrestrial lava flows correlate with 
reversed deep-sea sediments of the same age. It is clear 
that most reverse magnetization is acquired during 
periods of reverse polarity in the Earth's magnetic field. 
One of the major discoveries in paleomagnetism is 
that stratigraphic successions of volcanic rocks and deep-
sea sediment cores can be divided into sections that show 
dominantly reverse and normal magnetizations. Polar-
ity intervals are defined as segments of time in which 
the magnetic field is dominantly reversed or dominantly 
normal. Using magnetic data from volcanic rocks and 
deep-sea sediments, the Geomagnetic Time Scale was 
formulated (Cox, 1969), extending to about 5 Ma (Fig-
ure 1.18). Although polarity intervals of short duration 
(< 50 000 years) cannot be resolved with K-Ar dating 
of volcanic rocks, they can be dated by other methods 
in deep-sea sediments, which contain a continuous (or 
nearly continuous) record of the Earth's magnetic his-
tory for the last 100-200 My. The last reversal in the 
magnetic field occurred about 20 ka (the Laschamp). 
Two types of polarity intervals are defined on the 
basis of their average duration: a polarity event or 
subchron (10^-10^ y) and a polarity epoch or chron 
(10^-10^ y). A polarity chron may contain several-to-
z 
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MATUYAMA 
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Cochiti 
Nunivak 
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GILBERT 
Figure J .18 The Geomagnetic Time Scale for the last 5 
My. Grey pattern, normal polarity; white, reversed polarity. 
many polarity subchrons and can be dominantly normal 
(e.g., the Brunhes), dominantly reversed (e.g., the 
Matuyama), or mixed (Figure 1.18). Larger intervals 
(10 -^10*^ y) with few if any reversals are known as 
superchrons. Based on the distribution of oceanic mag-
netic anomalies, it is possible to extrapolate the Geo-
magnetic Time Scale to more than 100 Ma. Independent 
testing of this extrapolation from dated basalts indicates 
the predicted time scale is correct to within a few per-
centage points to at least 10 Ma. Results suggest that 
over the last 80 My the average length of polarity 
subchrons has decreased with time. Reversals in the 
Earth's field are documented throughout the Phanerozoic, 
although the Geomagnetic Time Scale cannot be continu-
ously extrapolated beyond about 200 Ma, the age of the 
oldest oceanic crust. Reversals, however, have been 
indentified in rocks as old as 3.5 Ga. 
The percentage of normal and reverse magnetization 
for any increment of time has also varied with time. 
The Mesozoic is characterized by dominantly normal 
polarities while the Paleozoic is chiefly reversed (Figure 
Plate tectonics 19 
100 
CEN I CRET I JUR ITRI IPERI CARB | D E V | S I L | ORD I CAM 
Number of Studies 
Figure 1.19 Distribution of 
magnetic reversals during the 
Phanerozoic averaged over 50 My 
intervals. Also shown are the 
Cretaceous (CN) and Permian-
Carboniferous (PCR) superchrons. 
Modified after Piper (1987). 
1.19). Periodic variations are suggested by the data at 
about 300, 110 and 60 Ma (Irving and Pullaiah, 1976). 
Statistical analysis of reversals in the magnetic field 
indicate a strong periodicity at about 30 My. Two major 
superchrons are identified in the last 350 My. These are 
the Cretaceous normal (CN) and Permian-Carboniferous 
reversed (PCR) superchrons (Figure 1.19). Statistical 
analysis of the youngest and best-defined part of the 
Geomagnetic Time Scale (< 185 Ma) shows an almost 
linear decrease in the frequency of reversals to the Cre-
taceous, reaching zero in the CN superchron. The inver-
sion frequency appears to have reached a maximun about 
10 Ma and has been declining to the present. Causes of 
changes in reversal frequency are generally attributed to 
changes in the relief and/or electrical conductivity along 
the core-mantle boundary. Both of these parameters are 
temperature-dependent and require long-term cyclical 
changes in the temperature at the base of the mantle. 
This, in turn, implies that heat transfer from the lower 
mantle is episodic. A possible source of episodic heat 
loss from the core is latent heat released as the inner 
core grows by episodic crystallization of iron. 
Vine and Matthews (1963) were the first to show that 
linear magnetic anomaly patterns on the ocean floors 
correlate with reversed and normal polarity intervals in 
the Geomagnetic Time Scale. This correlation is shown 
for a segment of the East Pacific rise in Figure 1.20. The 
correlations with polarity intervals are indicated at the 
bottom of the figure. A model profile for a half spread-
ing rate of 4.4 cm/y is also shown and is very similar to 
the observed profile. Both the Jaramillo and Olduvai 
subchrons produce sizeable magnetic anomalies in the 
Matuyama chron. The Kaena and Mammoth subchrons 
in the Gauss chron are not resolved, however. The lower 
limit of resolution of magnetic subchrons in anomaly 
profiles with current methods is about 30 000 years. Al-
though the distribution of magnetic anomalies seems to 
correlate well with polarity intervals, the amplitudes of 
anomalies can vary significantly between individual pro-
files. Such variation reflects, in part, inhomogeneous dis-
tribution of magnetite in oceanic basalts. Results suggest 
that most of the magnetization resides in the upper 0.5 
km of basalts in the oceanic crust (Harrison, 1987). Less 
than twenty per cent of the magnetization probably occurs 
below 0.5 km depth. 
100 100 km 
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Figure 1.20 Observed and model magnetic profiles across 
the East Pacific rise at 51 ° S latitude and corresponding 
correlation with magnetic polarity intervals for a spreading 
rate of 4.4 cm/y. After Vine (1966). 
Several detailed studies across magnetic reversals in 
stratigraphic successions provide information on the tim-
ing and details of the magnetic field behaviour during 
reversals (Bogue and Merrill, 1992). In those few sec-
tions where magnetic reversals are well dated, they oc-
cur over time intervals of as short as 1000 years and as 
long as 10 000 years. The best estimates seem to be near 
4000 years for the duration of a reversal. One record for 
a section across Tertiary basalt flows at Steens Moun-
tain in Oregon (-15 Ma) is shown in Figure 1.21. During 
this reversal, the intensity drops significantly and rapid 
and irregular changes in inclination and declination oc-
cur. In general, the field intensity drops 10-20 per cent 
during a reversal and suggests that a decrease in the 
dipole field precedes a reversal (Bogue and Merrill, 
1992). The actual intensity drop during a transition is 
latitude-dependent. Perhaps the most striking observa-
tion from the paleointensity record of the magnetic field