Download Magnetization of the oceanic crust

Geophys. J. R . astr. SOC. (1976) 41, 257-283
Magnetization of the oceanic crust
C .G. A. Harrison
University of Miami, Rosenstiel School of Marine and
Atmospheric Science, 4600 Rickenbacker Causeway, Miami, Florida 33149, USA
Received 1976 June 16; in original form 1976 February 13
Summary. The magnetization of the oceanic crust can be studied both
directly, by looking at the magnetization of rocks collected from the oceans,
and also indirectly, by looking at the magnetic anomalies. These methods are
discussed. Magnetic measurements on basement samples collected by the
Deep Sea Drilling Project (DSDP) show that these samples have lower magnetizations than dredged samples. The DSDP results suggest that the layer causing
the magnetic anomalies has to be much thicker than the normally assumed
thickness of 0.5 km. A theoretical power spectrum for lineated magnetic
anomalies is developed. It was assumed that the anomalies were caused by
normally- and reversely-magnetized bands of magnetic material formed during
periods of normal and reversed magnetic field, the lengths of the times of
constant field polarity being assumed to be exponentially distributed. Comparison of this theoretical power spectrum with published power spectra
also suggests that the magnetic anomaly sources are more deeply located
within the oceanic crust. The rock types which may contribute to the magnetic anomalies are discussed.
Introduction
It is known that the oceanic crust has to have a magnetization sufficient to cause the
lineated magnetic anomalies seen in many parts of the ocean basins. The origin of these
anomalies is to some extent unclear, in that the typical thicknesses of magnetized material
used to explain the anomalies require intensities of magnetization greater than those found
in dredge hauls of basement rocks and in samples of basalt recovered by the DSDP. This was
first pointed out by Lowrie (1 974).
In this paper I shall examine the data pertaining to the amount of magnetized material
necessary to explain the magnetic anomalies, the information relating to the intensity of
magnetization of samples obtamed from the ocean floor, and the power spectra of the
magnetic anomalies, in order to determine the possible sources for such anomalies.
9
258
C G. A . Harrison
1 Thickness of the magnetized layer
It has become common practice to assume that the magnetized layer is about 500 m thick
and then to adjust the intensity of magnetization in this layer to give the required amplitude
of magnetic anomalies. The origin of using such a thin layer is the work of Talwani, Windisch
& Langseth (1971) and Atwater & Mudie (1973). In order to determine the magnetization
of the oceanic crust, Talwani et al. obtained magnetic field measurements along lines parallel
to the ridge crest in the region of the Reykjanes Ridge. They found that the magnetic field
was highly correlated with the topography in this region, magnetic highs being observed over
topographic highs when the magnetization was normal, and magnetic lows being observed
over topographic highs when the magnetization was reversed. This is just what would be
expected at these high latitudes, where the magnetization and the magnetic field are both
predominantly vertical.
In order to explain the observed amplitude of the magnetic anomalies along the lines
running parallel to the ridge crest, Talwani et al. had to use certain values of intensity of
magnetization of the topographic relief. They found that in the central region of normal
magnetization, the intensity had to be as large as 30 Am-'. In a region which was 75 km
from the axis of spreading, and therefore about 7 Myr old, a value of 12 Am-' had to be
used. The assumption was made for these calculations by Talwani et al. that the topography
was two-dimensional, in other words that it continued without change perpendicular to the
profile. This is of course not in general true, and in fact there are many places on the ridge
crest where the small-scale topographic features are aligned parallel to the crest axis. If a
three-dimensional set of sources had been used instead of a two-dimensional set of sources,
the magnetization would have had to have been greater.
It should be pointed out here that Talwani et al. (1971) used a flat bottom for their
models of profiles running parallel to the ridge crest.
Atwater & Mudie (1973) have also obtained estimates of the intensity of magnetization
from looking at topographic features on the Gorda rise in the Northeast Pacific. They were
using observations of the magnetic field made close to the ocean floor, and found, like
Talwani et al. (1971), that in regions where they suspected normal magnetizations the
magnetic field was positively correlated with topography, whereas in regions where they
suspected reversed magnetization, there was a negative correlation between these two
parameters. To allow them to perform calculations to determine the intensity of magnetization, Atwater & Mudie (1973) had to make assumptions concerning the nature of the
magnetized layer. In some cases, where they believed the features to be constructional
features, built up by volcanic activity, they assumed that the lower layer of the magnetized
body was horizontal. In other cases where they believed that the topographic features were
produced by faulting, they assumed that the lower surface of the magnetized layer had the
same shape as the upper surface, or in other words, that the layer was of constant thickness.
They state that the shape of the lower surface is unimportant in determining the magnetic
anomalies, as these anomalies were observed very close to the bottom, and that it is almost
entirely the upper surface which governs the shape of the anomalies. This cannot be entirely
true, as if the magnetized layer were much thinner than the layer used by Atwater & Mudie,
the magnetizations would have to be greater in order to give the same amplitude of magnetic
anomalies.
The calculations of Atwater & Mudie (1973) assume linearity of the features which they
believe to be caused by faulting. As they state, if these features are in fact three-dimensional,
then the intensities of magnetization which they calculated are too small. They presented in
all 28 determinations of the intensity of magnetization, and these determinations show no
significant change according to their age. The mean value is 8.04 Am-' with a standard
Magnetization of the oceanic crust
259
deviation of 2.07 Am-'. The age of floor for which these intensities were calculated was
between 0 and 5 Ma.
One peculiar feature about the data from the Gorda rise is that there is this correlation
between topography and magnetic field. For instance Luyendyk (1969) showed that there is
very little correlation between topographic profiles and magnetic profiles obtained in the
Northeast Pacific over crust about 32 Myr old. Luyendyk assumed that the source of the
near-bottom profiles was in layer 2. However, this assumption was challenged by Peter
(1970) who pointed out that even if the near-bottom magnetic anomalies had their source
in layer 2 there was no reason to believe that the sea floor spreading anomalies observed at
the ocean surface also had their source in the same place. Klitgord (1974) has suggested that
some, but not all, of the short wavelength variation in the near bottom magnetic signal is
caused by topography. He also showed that many small-scale topographic features have little
or no magnetic signature. Many of the magnetic anomalies shown by Van den Akker,
Harrison & Mudie (1970) show no correlation with topographic features. Larson et 41.
(1974) concluded that some near-bottom anomalies were caused by topographic features,
but that some had to be caused by intensity variations.
On the basis of these high values of magnetization intensity, many people including
Atwater & Mudie (1973) and Talwani et al. have used rather thin layers of magnetized
material to model the sea floor spreading magnetic anomalies seen in many ocean basins.
These model studies will be discussed below, but one feature of them needs to be mentioned
here. This is that a uniform intensity of magnetization with depth is assumed. If, however,
the topographic features have a different magnetization than that of the crust beneath them,
then this assumption is not valid. For instance, some of the features may be constructional
features, as recognized by Atwater & Mudie. If these constructional features are formed of
slightly different material than the average upper oceanic crust, then they may have a
significantly different magnetization intensity. Also, because they are constructional
features, they are likely to have cooled down faster than the normal oceanic crust, resulting
in finer-grained material and a higher magnetization intensity.
Even if, as Atwater & Mudie claim, most of the features are caused by faulting, their
method only determines the magnetization intensity of the topmost portion of layer 2, and
if layer 2 has a decrease of magnetization with depth, as found in hole 332A of the DSDP,
then the methods of Atwater & Mudie (1973), and Talwani et al. (1971) d o not determine
iverage values of magnetization.
2 Intensities of magnetization necessary to explain the magnetic anomalies
2.1
INVERSIONS
We shall now describe the calculations which have been done showing what intensities are
necessary to explain the magnetic anomalies seen in the ocean basins. The most reliable
calculations are those done by inverting the magnetic anomaly data in order to arrive at the
magnetization intensity data. Since there is never a unique solution for the intensity of
magnetization following such an inversion, certain assumptions have t o be made. These are
usually that the magnetization is either parallel or anti-parallel to the axial dipole field at the
time of formation of the crust, or for young crust, to the present axial field. The top of the
magnetic body is usually taken to be the surface of the oceanic crust, or the surface of the
hard rock immediately underlying the sediment cover. The bottom is taken at different
places by different people. The most popular locations for the bottom are either 500 m
below the top, or 1.5 km below the top, this latter figure representing the average thickness
260
C. G. A. Ham'son
of seismic layer 2. In earlier publications, the bottom has been put even deeper. In analysing
these results, we shall not attempt to consider results from crust younger than 1 or 2 Myr.
There is much evidence that very young crust has an anomalously high magnetization intensity, and that some form of decay of magnetization takes place within a few Myr.
Bott & Hutton (1970) have inverted magnetic anomalies over the Sheba ridge in order
to obtain magnetizations of the ocean floor. The intensities of magnetization for anomalies
4-5 at the right-hand end of the profie shown in their Fig. 4 average about 2 Am-', for a
thickness of about 2 km. Eight of the block intensities have been left out of the calculation
of average intensity, as they were between zones of different polarity and so probably
represent areas where both positively- and negatively-magnetized material exists (Matthews
& Bath 1967; Harrison 1968). Therefore the average intensities of these blocks do not
represent the intensities of magnetization.
Bott (1967) presents an inversion of magnetic anomalies on the Juan de Fuca ridge.
Analysis of the values of the intensity of magnetization for oceanic crust older than about
2 Myr shows an average value of about 4.7 A m-l for a thickness of 1.7 km. It appears that
this value probably represents the value of the vertical magnetization, as earlier in the paper
a magnetization which is in the plane perpendicular to the strike of the two-dimensional
feature is introduced. Since the Juan de Fuca ridge runs approximately north-south, this
component will be the vertical component. Correcting this value to give a component along
the axial dipole direction produces a magnetization of 5.2 Am-'.
Emilia & Bodvarsson (1 969) have also inverted two anomaly sequences. One profile was
the Eltanin 19N profile, and they analysed magnetizations out to the beginning of the
Gauss normal epoch. If we take the values of magnetization for the left-hand end of their
profile, between the beginning of the Gauss, and the beginning of the Olduvai event, allow
for a slight change in the zero of magnetization, remove the four values of magnetization
which are close to reversal boundaries, we obtain an average magnetization of about
7 Am-'. These authors have used axial dipole directions of magnetization, so this value
does not have to be corrected. Doing the same procedure on the left-hand end of profde
V 20SA, and leaving out only two low values, gives a mean intensity of about 3 A m-'.
Parker & Huestis (1974) have inverted magnetic anomalies obtained close to the sea
floor, and spanning the Gilsa event (Watkins 1972). The mean magnetization change on
either side of the Gilsa event is about 29 Am-', giving a mean magnetization of about
14 A m-'. The reason that this value is much higher than the values given previously is that
the thickness of the magnetized layer was assumed to be 0.5 km, following Talwani et d.
(1971) and Atwater & Mudie (1973). Parker & Huestis (1974) show, however, that the
intensity of magnetization varies almost inversely with the thickness of the magnetized
layer, down to a layer thickness of 1 km.
Klitgord (1974) has presented an analysis of magnetization changes across reversal
boundaries. He inverted magnetic field data from the deep tow of SIO taken over five
actively-spreading ridges, and then determined the average magnetization contrast through
reversal boundaries. One half of this value is then the average magnetization. His data show a
pronounced decrease in the magnetization contrast as older crust is encountered. However,
the data for crust between the ages of 4 and 5 Myr do not show this decrease, which happens
in crust 0-4 Ma old. He assumed that the magnetized layer was 0.5 km thick, and obtained
between 4 and 6 magnetization values between the two ages quoted, for five different ridge
segments. The averages of these values are given in Table 1.
Harrison & Mudie (unpublished data) inverted deep observations of the magnetic field
going across the older boundary of anomaly 14 in the Northeast Pacific at latitude 34" N.
The thickness of the magnetized blocks used in the inversion calculation was 4 km, which is
on data for intensity of magnetization
1970
arsson 1969
arsson 1969
is 1974
die, unpubl.
die, unpubl.
Age
(Ma)
7-10
2-3s
2-3s
2-3%
2
4-5
4-5
4-5
4-5
4-5
39
56
Thickness
(km)
Magnetization
(A m-')
Latitude
N)
e
Paleolatitude
(ON)
Equatorial
Mag"
(A m-')
2.0
1.7
2.0
2.0
0.5
0.5
0.5
0.5
0.5
0.5
4.0
1.4
2.08
5.17
7.12
3.08
14.62
17.26
11.89
9.45
6.82
4.04
1.55
5.38
15
47
- 52
- 28
-51
47
-51
41
21
3
34
35
15
47
-52
- 28
-51
47
-51
41
21
3
22
18
1.898
3.203
4.208
2.390
8.719
10.695
7.091
6.243
5.794
4.024
1.303
4.743
Equatorial
Mag" for
0.5 km thick
layer
1.592
10.890
16.832
9.560
8.719
10.695
7.091
6.243
5.794
4.024
10.424
13.280
Locat
Sheba
Juan
East P
Mid-A
Pacifi
Juan
Pacifi
Gorda
East P
Costa
North
North
262
C. G. A . Ham‘son
a lot thicker than most of the calculations already described. The change in magnetization
on going across the boundary was about 3 A m-’, giving an average magnetization of half this
amount. In further unpublished data, Harrison & Mudie inverted both surface observations
and deep observations of the magnetic field across anomaly 22 in the Northeast Pacific at
latitude 35” N. The average magnetization contrast on either side of this anomaly was about
11 A m-l for a thickness of 1.4 km, giving an average magnetization of half this value.
All these results are listed in Table 1. Since all of the results have been presented assuming
that the direction of magnetization is along an axial dipole magnetic field, it is possible to
correct for the latitudinal effect of the magnitude of the dipole field. This is done by
” ~ h is the latitude, to give a value corrected to the equator.
dividing by (1 + 3 ~ i n ~ X )where
The polar value will be just twice the equatorial value. In the case of the last two values, a
paleolatitude was used instead of the present latitude, as these two values are for old crust
which has moved northwards significantly since it was formed (Francheteau et af. 1970;
Harrison et af. 1975).
The data discussed above and presented in Table 1 give 12 estimates for magnetization
of the oceanic crust, which, however, vary according to the model used. The main variation
is, of course, in the thickness of the magnetic layer. We can investigate the variation of
magnetization intensity with both thickness and age by doing a multiple correlation analysis
of these three variables, using the magnetizations reduced to equatorial values as the
dependent variable. Since the magnetization is approximately correlated with the reciprocal
of the thickness, we use this reciprocal as an independent variable.
The equation representing the variation of magnetization with thickness and age is then
2.926
I = t 0.009 T + 1.245
d
where I is measured in A m-l, d in km, and Tin Ma.
D, K M
4.0 2.0
(D =
1.0
0.5
I/D
Thickness in K M )
Figure 1. Magnetization as a function of thickness. Small dots refer to individual values. Large dots refer
to mean values for each thickness. If the large dot is solid, it is also the position for a small dot. The
straight line through the origin also goes through the mean value of magnetization reduced to a common
thickness of 0.5 km (see text).
Magnetization of the oceanic crust
263
The multiple correlation coefficient is 0.7925 and is significant at almost 1 per cent. The
partial correlation coefficient for the regression of I caused by l/d is 0.7704 and is
significant at 1 per cent. The partial correlation coefficient for the regression of I caused by
T is almost zero (-0.0810) and is not significant, as could have been deduced by the very
small value of the multiplier of Tin the above equation. The correlation between magnetization and thickness is quite striking. Fig. 1 shows a plot of magnetization versus the reciprocal of thickness, average magnetizations being shown where there are several observations
using the same thickness of the magnetized layer.
These magnetizations can be reduced to magnetizations for a common thickness by the
following crude method. If there is a discontinuity in vertical magnetization from +Ito -I
along a vertical plane in an infinite sheet of magnetized material, then the maximum vertical
induction produced close to this discontinuity is given by
= 41 [tan-'
-
tan-'
&]
where h is the depth to the top of the sheet and d is the sheet thickness.
The same formula is true for discontinuities in horizontal component of magnetization
and horizontal field. Therefore, to obtain the same maximum field from two different
thicknesses d and d : the intensities of magnetization are in the ratio
-
Id'
- -
tan-'
[24i@T2-3 3
dl
For small values of d and d' with respect to h , this formula shows that the intensity of
magnetization is directly correlated with the reciprocal of thickness. As the thickness
becomes a significant part of the depth to the top of the slab, the magnetization does not
decrease as fast. In other words, doubling the thickness from h / 2 to h causes a change of
magnetization by a factor of 0.593. This non-linearity may become important for many of
the observations of the field done close to the ocean bottom, as in this case h is very small.
However, this relationship only applies to the maximum or minimum anomaly on either
side of the discontinuity. If we consider the anomalous field produced at the centre of a
block which is several tens of kilometres wide, it can be shown that this anomalous field is
almost dependent on the thickness of the block, for thicknesses less than that of the oceanic
crust, since the field depends on the angle subtended at the point of observation by the
edges of the block. Parker & Heustis (1974) have pointed out that even in the case where
they were looking at anomalies about 0.5 km above the top of the basement, on changing
the thickness from 0.1 to 1.0 km, the intensity fell almost by a factor of 10. Therefore, in
adjusting to a common thickness we have simply assumed that the intensity is directly proportional to the reciprocal thickness. These values are given in Table 1 . The mean of the
12 values is 9.26 Am-' with a standard error of 1.01 A m-I, these figures being for a
thickness of 0.5 km.
2.2
DIRECT METHODS
There are other estimates of the intensity of magnetization necessary to explain the magnetic anomalies. These come from the forward method which is used to create theoretical
264
C. G. A . Ham'son
model anomalies for any given time period, and for the parameters of spreading rate and
direction applicable to the portion of the ocean under consideration. By comparing the
computed and observed anomaly amplitudes it is possible to get a rough estimate of the
actual magnetization necessary to give the correct amplitude for the given layer thickness.
Some of these results will be discussed here, and are summarized in Table 2.
Table 2. Comparison of observed and simulated anomalies
0.5 km
Reference
Age
(Ma)
1-9
Talwanietal. (1971)
Herron (1972)
1-70
McKenzietkSclater (1971) 1-70
Vine & Hess (1970)
60-70
Pitman&Talwani (1972)
1-70
Thickness
(km)
Magneti- Latization
tude
(A m - ' ) (ON)
0.4
0.5
2.0
1.8
0.5
12.0
10.0
2.5
5.0
10-20
Paleolatitude
(ON)
thick
layer
(A m-')
61
0,-45
20, -50
61
9.6
-
10.0
-
10.0
so
30
15,ss
-
18.0
10-20
Location
Reykjanes Ridge
South-east Pacific
Indian Ocean
North-east Pacific
North Atlantic
Talwani et af. (1971) have used a magnetization of 12 A m-' for crust older than 1 Ma
to explain fairly well the amplitudes of the magnetic anomalies seen over the Reykjanes
ridge with a 0.4-km thick magnetized layer. Although it is not specifically stated in their
paper, 1 have assumed that this magnetization is the magnetization along the axial field.
Herron (1972) has presented many model anomalies for the South-east Pacific including
the East Pacific rise, the Galapagos spreading centre, and the Chile rise. The general impression is that her choice of thickness and intensity gives approximately the correct
amplitude of anomalies. In some cases the theoretical anomalies are larger, and in some
cases they are smaller, than the observed anomalies.
McKenzie & Sclater (1971) generated synthetic magnetic anomalies in the Indian Ocean,
using thicknesses of 2 km for most of the region they considered. Instead of using intensities
of magnetization, McKenzie & Sclater used susceptibilities, the value chosen being 0.01. By
doing this they automatically allow for the latitude effect, and the value of the equatorial
magnetization is 3.1 A m-'. The general impression of a comparison between the observed
and computed anomalies is that the computed anomalies are never smaller than the observed
anomalies, but are sometimes up to 1.5 times greater. Therefore a better magnetization
would be about 2.5 A m-'.
Vine & Hess (1970) have presented data from the Great Magnetic Bight in the North-east
Pacific. The comparison between their simulated and observed profiles along the northsouth direction shows good agreement of amplitude, and therefore we may assume that the
magnetization of 5 A m-' for a layer of 1.8 km thick is a good estimate. The observed
east-west profiles show a smaller amplitude than the simulated profiles. This is because the
Pacific floor was further south at the time of formation than it is today, and this makes a
difference to the north-south trending anomalies, but not very much, except in shape, to
the east-west trending anomalies. The ratio of the amplitude of these sets of anomalies
has been used by Vine (1968) to give the northward movement of the Pacific since the time
of their formation.
Pitman & Talwani (1972) have studied magnetic anomalies in the North Atlantic, and
have compared them with anomalies simulated from a layer 0.5 km thick, with an intensity
of 15 A m-' along the direction of the axial dipole field. There is a great deal of variation
between the sizes of the observed anomalies compared with the sizes of the computed
anomalies. This makes it difficult to judge whether the magnetization assumed is in general
correct or not. I have assumed a possible error of one third.
265
Magnetization of the oceanic crust
It is fairly obvious from Table 2 that the equatorial magnetization for a thckness of
0.5 km is very similar to the value determined from the inversions shown in Table 1, or
about 10 A m-'. This value, and a thickness of 0.5 km, has been used in numerous other
modelling studies, and seems to be the standard for most theories of the origin of the
magnetic anomalies.
2.3
SEAMOUNTS
Inversion schemes are available to deal with the magnetic anomalies produced by threedimensional bodies such as seamounts. These inversion schemes have been used to calculate
the directions of magnetization in seamounts and hence to estimate a paleomagnetic pole
(Francheteau et al. 1970). Intensities of magnetization are also calculated at the same time.
Results from 30 seamounts in the North Pacific (Harrison et al. 1975) give a mean intensity
of magnetization of 6.77 A m-' with a standard error of 0.63 A m-'. These 30 seamounts
were ones which had a moderately uniform magnetization. Some of the seamounts appeared
to have non-magnetic tops formed of hyaloclastites (Harrison 1971), and this has been
allowed for in many of the calculations, but not all. Other results from 35 Pacific seamounts
give less good agreement between observed and computed anomalies, but the mean magnetization of these is slightly higher than for the 30 good results.
Younger seamounts from the Pacific give somewhat smaller average intensities of
magnetization. But in this case, we must remember that the seamounts are likely to have
formed during several different polarities of the Earth's magnetic field, which probably
explains the somewhat lower average intensities, and also the fact that most of these Tertiary
seamounts do not give very good agreement between observed and computed anomalies
(Francheteau et al. 1970).
Five Gulf of Guinea seamounts give a rather high mean intensity (9.28 A m-'; Harrison
1970), as d o seven of the Kelvin seamounts (10.93 Am-'; Richards, Vacquier & Van
Voorhis 1967). All these results are summarized in Table 3. Since these results are from
contrasts of magnetization between basalt and water, there is very little error in determining
the intensity of magnetization. We d o not have the problem of the lineated magnetic
anomalies, where the intensity is essentially unknown until we know the depth and thickness of the layer causing the anomalies.
Table 3. Seamount magnetization
Seamount Group
North Pacific (Good)
North Pacific (Poor)
North Pacific
North Pacific
Gulf of Guinea
Kelvin
Number
Age
Mean Equatorial
Intensity (A m-') Standard error
30
Cretaceous
Cretaceous?
Tertiary
Upper Tertiary
Cretaceous?
Cretaceous?
6.41
7.04
4.23
3.86
9.28
10.93
35
10
16
5
7
References:
1.
2.
3.
4.
Harrison et af. (1975)
Francheteau et al. (1970)
Harrison (1970, 1971)
Richards etal. (1967)
0.60
1.10
1.04
0.58
1.54
0.77
Reference
1
1, 2
2
2
3
4
266
C. G. A. Ham'son
Discussion: magnetizations determined by indirect methods
Presentation of the foregoing data suggests several things. Seamounts have quite high values
of magnetization, of around 4-10 A m-', with an overall average value of 6.5 A m-', when
reduced to equatorial values. These seamounts are composed of rapidly-cooled piles of
extrusive volcanics, some of which have reacted with sea water to form hyaloclastites. Hydoclastites have a very low magnetization, insufficient to cause pronounced magnetic anomalies
(Harrison & Ball 1974). In some cases, the presence of hyaloclastites in the seamounts has
been allowed for by making calculations in which the top portion of the seamounts has been
made non-magnetic. We have rather little evidence as to how frequently extrusive flows
which build up seamounts turn into hyaloclastites. It has been assumed previously (Harrison
1971) that the magnetic anomaly over seamounts is caused by the magnetization of the
feeder dykes and basalts which intrude into earlier portions of the volcanic pile, these
features being protected from sea water and hence not changed into hyaloclastite. However,
this means also that they would be cooled more slowly than if they were in contact with
sea water. The very high magnetizations of the seamounts suggest in contrast that the
material was rapidly cooled. These inconsistencies cannot be answered at this time. However,
it seems fairly clear that the magnetizations shown in Table 3 are probably minimum
magnetizations for the strongly magnetic portion of the seamounts.
Tables I and 2 suggest that if the layer causing the magnetic anomalies is the commonly
assumed thickness of 0.5 km,then the magnetization produced by the equatorial field would
have to be in the region of 10 A m-', in order to produce the required amplitude of the
anomalies. It should also be the average magnetization of the top of the oceanic crust, as
sampled by the Glomar Challenger during the DSDP and the International Phase'of Ocean
Drilling. The following section is a compilation of the intensities of magnetization obtained
from samples of the oceanic crust, and was done in order to compare the results of
magnetization intensities obtained by the indirect method with those obtained by the
direct method.
Direct measurements of the oceanic crust
Lowrie (1974) was the first person to compile magnetization results from the DSDP and
from dredge hauls. He showed that the mean NRM intensity of dredged rocks was almost
twice as great as the mean intensity of drilled rocks. I believe that the correct mean value to
take in these circumstances is the arithmetic mean value. Although most suites of rocks give
magnetic parameters such as NRM intensity, or susceptibility, which have a log normal distribution, this does not mean that the mean has to be the geometric mean, which is that used
by Lowrie. For instance, if layer 2 were composed of equal amounts of fresh basalt of
intensity 5 A m-' and metamorphosed basalt of intensity 0.5 A m-', and if our method of
collection brought up a representative suite (i.e. equal numbers of each rock type), then the
arithmetic mean intensity of magnetization would be 2.75 A m-', the geometric mean would
be 1.58 A m-', but the correct value to use in determining the anomalous magnetic field
would be the arithmetic mean value, which is always higher than the geometric mean. Study
of the papers quoted by Lowrie suggests that the arithmetic mean value of NRM from
dredged sample is considerably higher, or about 10 A m-' compared with his geometric mean
of about 5.5 A m-'. This value includes all types of basalts (fresh and weathered) except
metamorphosed basalts, which have very much lower magnetization intensities than even
highly weathered basalts (Fox & Opdyke 1973). That this difference between the geometric
Magnetization of the oceanic crust
267
and arithmetic mean is not unreasonable can be seen by calculating the two means for the
62 samples of unmetamorphosed basalt given in Fox & Opdyke (1973). The arithmetic
mean is 5.3 A m-' and the geometric mean is 2.1 A m-'.
I have calculated mean values of magnetization of basement rocks in 50 DSDP holes
(Fig. 2), in order to update the information provided by Lowrie (1974). In order to be
consistent with the results presented for the inverted magnetic anomaly profiles, the mean
magnetizations in each core have been corrected to give equatorial values by dividing by
(1 + 3 ~ i n ~ h )where
' ' ~ h is the paleolatitude of the site when it was formed at the midoceanic ridge system. The paleolatitude is estimated from paleomagnetic results for adjoining
continents given by McElhinny (1973). In the case of the Pacific, there is no adjoining continent and in this case the paleolatitudes are estimated from the results of seamount surveys
in Francheteau et al. (1970) and Harrison et al. (1975). In the few cases where the inclinations of the basalts were both positive and negative, and greater than 30°, the mean intensity
was established by treating one group as having negative intensities, as the most reasonable
assumption concerning mixed inclinations is that there are both normally- and reverselymagnetized rocks present. Very little difference would have ensued if this method had not
been used.
These results are given in Table 4 and Fig. 3. The mean of all 50 values is 2.41 k 0.27 (se)
A m-'. Many of the holes were drilled in areas in which the lineated magnetic anomalies have
not been recognized or correlated. If the mean value for the 21 cores which fall on areas of
lineated magnetic anomalies (identified by using the map of Pitman, Larson & Herron 1974)
is calculated, it is found to be only 1.61 ? 0.35 (se) A m-', or a factor of 5.75 less than the
average value of a 0.5-km thick layer from Table 1. It could be argued that the value derived
from the DSDP holes is much lower than the upper portion of layer 2, either because of
weathering of the uppermost layers, or because of some other unknown effect which
increases the intensity of magnetization with depth. This possible effect has been studied in
two different ways. The first way is to study the magnetizations observed in DSDP holes as a
function of the age of the basement. Although some of the samples of hard rock have
penetrated what are thought to be sills (Lowrie 1974), the age of these sills is unlikely to be
many millions of years younger than the true basement age. Hence the basement ages are
assumed to be the ages of the hard rocks sampled by the DSDP. Fig. 4 shows a plot of mean
magnetization with respect to age. The least-squares regression line for all 50 cores is shown,
and has a correlation coefficient of 0.353 which is significant at the 5 per cent level. So these
data suggest that any weathering effect, which would tend to reduce the magnetization of
the older rocks more than the younger rocks, is not an important factor in determining the
mean magnetization of the rocks.
The second way is to determine directly the vertical variation of magnetization in individual holes, where such holes have penetrated significant depths into the oceanic basement.
If weathering has been more important in reducing the magnetization of the surface rocks
than the deeper rocks, then such an effect should be seen in vertical variations of magnetization seen in individual holes. The most important holes are those which have penetrated
deepest into the oceanic crust, which were drilled on leg 37. The magnetic measurements
are described by Scientific Party (1975). In hole 332A, which penetrated 333 m of basalt
interlayered with lithified ooze, above which was 104 m of ooze, the magnetization of the
basalt clearly showed a decrease of magnetization with depth (Fig. 5). This decrease of
magnetization is most clearly seen in the 100-m average values of intensity of magnetization. These 100-m averages were obtained from the average values of individual 9.5-m cores,
although the same picture would have been obtained if each measurement had been given
equal weight. The 100-m averages given 4.0Am-' for the top l o o m , followed by
I
I
I
I
Location of DSDP holes which have recovered basalt whose magnetic properties have been studied. The asterisks are holes which a
agnetic anomalies, and the squares are holes in areas where heated magnetic anomalies have not yet been identified.
269
Magnetization of the oceanic crust
Table 4. Magnetization intensity in oceanic basement drill holes
Age
(Ma)
23
85
39
21
25
53
12
15
23
40
11
8
109
101
96
65
84
87
79
73
84
74
50
30
67
53
44
55
78
62
58
131
84
22
46
38
101
106
109
102
158
21
48
40
63
106
23
23
24
40
Hole
Experimental
Mohole
10 DSDP
14
15
18
19
36
54
57
77
83
84
136
137
138
141
146
150
151
152
153
163
183
191
192
220
22 1
223
239
24 5
24 8
24 9
250
25 1
253
254
256
257
259
260
26 1
279A
280A
282
283
317A
319
319A
320B
321
No. of
sample
Paleolatitude
(ON)
Mean equat.
magnetization
(A m-*)
On
lineation
Reference
J
1
J
2, 3
2, 3
2, 3
2, 3
2, 3
3
3
3
3
3
3
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
4
23
24
4.400
20
9
3
7
14
7
6
15
8
3
9
2
2
4
4
12
3
2
3
3
1
2
5
9
17
6
7
2
2
2
2
6
4
1
9
17
30
9
3
9
4
11
13
2
34
2
13
2
8
17
-27
- 29
- 32
- 29
39
16
9
-11
0
3
39
31
31
14
22
22
23
23
21
-13
32
61
28
- 25
- 18
4
- 26
- 36
- 39
- 45
-52
- 37
-50
-57
-50
- 60
-63
- 44
- 40
- 65
- 67
-58
- 69
- 39
-21
-21
-11
- 13
1.071
0.527
0.077
0.884
1.149
0.554
0.839
3.089
1.804
1.500
6.97 1
7.573
3.739
2.866
3.832
2.965
4.264
3.230
3.418
5.157
1.118
1.842
0.551
1.896
2.299
1.098
4.754
1.277
3.630
3.438
2.384
6.324
3.105
0.530
0.895
3.504
1.649
1.886
7.645
1.787
1.464
0.068
0.529
0.378
3.104
0.05 1
0.974
0.333
2.006
J
J
J
J
J
X
X
J
J
J
X
X
X
X
X
X
X
X
X
J
J
X
J
X
X
X
J
J
X
X
X
X
X
X
X
X
X
X
X
J
X
X
J
X
J
J
J
J
4
4
4
4
4
4
4
4
4
4
4
5
5
5
5
References:
1. COX& Doell (1962)
2. Lowrie et al. (1973a)
3. Lowrie er al. (1973b) 4. See results in Initial
Reports of the DSDP for the appropriate leg. 5. Tarasiewicz, Tarasiewicz & Harrison (1976)
270
C. G. A . Harrison
12
a
W
m
3
z 8
4
0
16 1
1.61
0
2
4
6
8
A m-I
Figure 3. Histogram of mean hole intensities for 50 DSDP holes. The values have been reduced to those
appropriate for the equator as described in the text. The solid portions are for the holes drilled in areas
where h e a t e d magnetic anomalies have been seen. The lower histogram is plotted on a linear scale and
the arithmetic means are indicated above, for the two sets of data. The upper histogram is plotted on a
logarithmic scale, and the geometric mean values are indicated.
e,ol0
= LINEATIONS
NO LINEATIONS
.
E 6.0
a
* . O
0-
0
(D
0
0
0
.
I
.
I
Age, MY
Figure 4. The mean magnetization intensity from 50 DSDP holes plotted against the age of the basement
The magnetizations have been reduced to equatorial values, as described in the text. The line is the leastsquares regression line for all 50 points, with age as the independent variable.
27 1
Magnetization of the oceanic crust
DISTANCE FROM TOP OF EASEMENT, M.
0
100
200
400
300
500
10
8-
332 B
-
-
6-
I
E
a
4-
2Y
04
0
10
20
30
40
50
60
DISTANCE FROM TOP OF BASEMENT. CORE LENGTH OF 9.5 M
Figure 5. Core average intensities of magnetization from 2 DSDP holes which penetrated several hundred
metres into the basement. The straight lines are the least-squares regression lines through these data,
assuming that the depth is the independent variable.
3.2 Am-' and 2.3 A m-' for the next two 100-m sections, with an average of 1.6 A m-'
for the bottom 33-m of basalt. The average values of magnetization for the layers are in
fact lower than this because of the intercalated sediments which are of course much more
weakly magnetized. Parenthetically, this hole also showed both positive and negative values
of the inclination, which, if interpreted as normally- and reversely-magnetized material,
showed that at this one position the magnetic effect of the top 333-m of basement had
almost n o effect on the magnetic anomaly.
In hole 332B, which penetrated 589 m into the basement, there is a tendency for an
increase in magnetization with depth (Fig. 5). In all, 41 out of the 47 cores recovered
material for which there are measurements of magnetization. If the core numbers are taken
to indicate depth, then the correlation coefficient between core number (or depth) and
mean core magnetization is 0.0537 which is not significant at the 5 per cent level. The bestfitting line through the data points has a slope of 0.0149 A m-' per core length. So in the
46 core intervals representing 589-m of basement, the average magnetization rises by an
amount of 0.685 A m-', an insignificant amount.
An interesting feature of hole 332B (and other holes as well) is that the measured seismic
velocity on the recovered samples of basalt is much higher than results from seismic
refraction experiments would indicate. Hyndman et al. (1976) have suggested that this is due
to intercalation of rubble and sediments into the solid basalt within the top few hundred
metres of the crust sampled at this drill site. Based on the relative drilling effort, they found
that of the total thickness of 637 m of layer 2 sampled, about 44 per cent was probably
low-velocity sediment and rubble. Therefore the measured magnetizations on material
recovered from this site, almost all of which was solid basalt, should be reduced by almost a
factor of 2 to determine the average magnetization of the layer.
Hole 333A penetrated 312 m into hard rock. In this case there was a tendency for the
samples lower in the section t o be more strongly magnetized than those higher i n the
272
C. C. A . Harrison
section, although the correlation coefficient of magnetization as a function of depth is not
significant even at the 10 per cent level. Both hole 334 (penetration of 48 m into basement)
and hole 335 (penetration of 114 m into basement) showed a tendency for magnetization to
decrease with depth, although again the correlation coefficients are not significant. All these
data are summarized in Table 5. The correlation coefficients were all calculated by taking
the core arithmetic average magnetizations as a function of depth to the top of the cores.
The gradient of magnetization with depth is expressed in A m-' over a depth of 500 m.
There seems little justification in these data for assuming that there will be in general an
increase in magnetization with depth, at least over the top few hundreds of metres.
Table 5. Change of magnetization with depth in samples from Leg 37 of DSDP
Basem ent
penetration
Hole No.
(m)
Correlation
coefficient
(r)
332A
332B
333A
334
335
333
582
312
48
114
-0.3764
0.3007
0.3366
- 0.6974
-0.3399
Significance
of r
(%)
Number of mean
magnetizations
available
5-10
5-10
> 10
>10
> 10
21
41
10
5
11
Slope
(A m-I /500m)
- 3.830
2.462
4.322
-36.158
-13.194
Other cores which have penetrated the oceanic crust also do not suggest any significant
increase of magnetization with depth. Cox & Doell (1962) found that in the experimental
Mohole core, there was a tendency for the magnetization to decrease with depth, whereas
in hole 317A (Jarrard, private communication) there is a significant increase with depth.
Therefore there seems to be no reason not to conclude that the NRM values measured in
the short DSDP drill holes represent average values of the magnetization in the upper layers
of the oceanic crust.
Viscous magnetization
NRM values only give the correct value to use in determining magnetic anomalies if these
values are for stable magnetization. If for instance the original TRM for lavas was 5 A m-'
and all lavas have acquired a viscous component of magnetization in the recent period of
normal magnetic field of magnitude 3 A m-I, then the normally-magnetized lavas would have
a magnetization of 8 A m-', whereas the reversely-magnetized lavas would have a magnetization of only 2 A m-'.
The magnetization contrast would still be 10 A m-I. (These calculations assume that TRM
and VRM are additive.) The average magnetization of these samples will still be 5 A m-', half
the magnetization contrast. If however the viscous magnetization acquired is greater than the
original NRM value, say 6 A m-', then the magnetizations of originally normally- and reverselymagnetized rocks will now be 1 1 and 1 A m-' respectively. The magnetization contrast will
still be 10 A m-', but the average intensity of magnetization will now be 6 A m-', leading the
unaware observer to think that the magnetization contrast should in fact be 12 Am-'.
In other words, if the viscous component is large enough to reverse the original direction of
magnetization of reversely-magnetized rocks, then the average magnetization calculated by
averaging the hole averages will be too large a value t o use when magnetic anomalies are
being considered. This may be an important effect in some holes, as h w r i e , Uvlie &
Opdyke (1973b) have shown that in some cores, the magnetization measured has the appearance. of being almost entirely viscous. Hence the average value given is probably on the
large side.
Magnetization of the oceanic crust
273
Discussion
Simple calculations have shown that assuming a constant magnetization vector within the
upper layers of the oceanic basement, we would have to take between about 2 and 3 km of
material in order t o produce the required intensity of magnetic anomalies observed,
depending on which average magnetization is used. Within the median valley of the midAtlantic ridge many rock types are dredged, including the extrusive basalts. These other rock
types are gabbros, metamorphosed basalts and gabbros including amphibolites, and various
ultramafic rocks. All these rock types have magnetizations lower than the extrusive basalts
(Fox & Opdyke 1973). The presence of these other rock types suggests that the extrusive
layer is indeed quite thin, as was found in hole 334 of the DSDP (Scientific Party 1975). The
possibility of sampling deep into the oceanic crust within the median valley is limited to the
thickness exposed along the largest exposed fault scarp, which is usually 500 m, and never
more than 1 km in extent. The fact that many other rock types are found in the median
valley suggests that the extrusive layer cannot be more than about 5-1000 m in thickness.
This extrusive layer may represent the upper portion of layer 2 , as described by Talwani
et al. (197 1) and by Poehls (1974) who discovered a seismic layer in the mid-Atlantic ridge
with an average thickness of about 1 km at the top of the oceanic crust, with a rather low
velocity of between 3 and 4 km s-l. This suggests that there have to be magnetic sources
deeper within the oceanic crust which are partially responsible for the production of the
magnetic anomalies.
A word should be said here about the seeming contradiction between the results of
Talwani ef al. (1971) and Atwater & Mudie (1973) on the one hand, and the results
described above on the other hand. As has been noted before, the thinness of the magnetic
layer derived in the above two papers was calculated assuming a constant intensity with
depth. If the topographic features have a greater magnetization than the material below,
then these calculations of thickness are invalid. There is good reason to suppose that the
topographic features are indeed more magnetic than the material below. This can be seen
from a comparison between the magnetization of dredged material and the magnetization
of DSDP material. As Lowrie (1974) has demonstrated, the dredged samples have a higher
magnetization than the DSDP samples. Part of this is no doubt due t o the fact that many
dredge stations have been made close to the axis of the ridge, where rock outcrops are
common. As has been pointed out by Irving, Robertson & Aumento (1970) rocks collected
close to the ridge crest have anomalously high magnetizations, and so a bias is put into the
mean magnetization of dredged samples as compared to DSDP samples, which never sample
close to the ridge crest because of the lack of sediment cover there.
But some results have been obtained from older material. For instance, Matthews (1961)
analysed 100 samples of highly altered vesicular basalt dredged from a 200-m high abyssal
hill in the North-east Atlantic where the ocean floor age was lower Tertiary. The median
magnetization was 5 A m-', which when corrected for latitude gives an equatorial value of
4 Am-'. Very few of the DSDP holes have average magnetizations greater than this value
(Fig. 3). Also the very high magnetizations obtained from seamounts suggests that topographic features can be strongly magnetized.
DSDP holes are frequently drilled into topographic lows, whereas dredged rocks usually
come from topographic highs. Therefore the supposition that the topographic highs are more
strongly magnetized than the material below agrees with the observation that dredged
samples are more strongly magnetized than DSDP samples, and with the supposition that the
methods of Talwani et al. (1971) and Atwater & Mudie (1973) for obtaining the thickness
of the magnetized layer will give values which are too small.
Studies of the Macquarie Island ophiolite complex also suggest that lower layers may
274
C. G. A . Harrison
be important sources of magnetic anomalies. Butler & Banerjee (1Y73) and Butler, Banejee
& Stout (1975) have suggested that the pillow lava sequence in this complex suffers progressive decay of magnetization with age. The underlying dyke swarm sequence has a similar
intensity of magnetization to the pillow lavas, but the magnetization is more stable in the
lower layer. Butler & Banerjee (1973) also suggest that gabbroic and ultramafic (serpentinite)
layers could also be important contributors to marine magnetic anomalies.
Power spectra for magnetic anomalies
In this section we shall derive a theoretical spectrum for marine magnetic anomalies in order
to compare these spectra with calculated spectra. This work was stimulated by a paper of
Blakely, Cox & Eufer (1973) and specifically by their analysis of a three-component aeromagnetic profile obtained in the North-east Pacific. The profrle was about 900 km long and
was obtained over oceanic crust dated between 46 and 65 Ma by the magnetic anomalies
present, using the time scale of Heirtzler et al. (1968). They presented a power spectrum of
this profile, and pointed out that this spectrum has two parts. The part at low wavenumbers
is that caused by the h e a t e d magnetic anomalies. The part at wavenumbers greater than
0.3 rad km-’ is caused by aircraft motion. They also pointed out that the slope of the plot
of the natural logarithm of power with respect to wavenumber could be used to determine
the depth to the source of the magnetic anomalies by employing the method of Spector &
Grant (1970). In this method, the slope is shown to be -22 at high wavenumbers, where z
is the depth to the sources. Since the portion of the power spectrum which is caused by
crustal magnetization variation is limited to wavenumbers less than 0.3 rad km-’ by noise
at higher wavenumbers, it is not possible to employ Spector & Grant’s method directly.
Instead, what we do is to determine a theoretical power spectrum for sea floor spreading
anomalies in order to see what parameters are necessary to match the shape of the observed
power spectrum.
We start by assuming that the reversals of field occur randomly with time such that the
lengths of polarity intervals are like an exponential distribution. This has been shown to be
approximately true for the interval under consideration by Blakely & Cox (1972) if the
shorter polarity events discussed by them are real. In this case, the power spectrum of the
magnetization can be written down. Rice (1954) has shown that the power spectrum of a
signal of amplitude & A in which the lengths of each period of constant amplitude are
exponentially distributed is given by
2pA2
Wf)= 7
p +7?f2’
where p is the rate of change (i.e. on average there are p changes of sign per unit time), and
f is the frequency.
If we assume that p is the rate of reversals per Myr, that k is the wavenumber in rad km-’
and that s is the spreading rate on km Myr-’ , then the power spectrum as a function of
wavenumber will be
Thus the source of magnetic field anomaly does not have a ‘white’ spectrum as is often
assumed.
Magnetization of the oceanic crust
275
If the magnetized layer is assumed to have a vertical magnetization, then we can calculate
the field produced by a surface of magnetic poles in the following way. Assume the sheet of
poles to be at a depth h and to vary along the y direction sinusoidally with wavenumber k.
Variation in the other horizontal direction is assumed zero as the anomalies are heated.
Then the vertical field at any pointy’ is given by
I_
2 A h sin (ky)dy
Fv =
0,-y’)’+h2
.
Therefore F,, = 2 A sin (ky‘) n exp (-kh),
where A is the amplitude of the magnetization variation which produces the sheet of poles.
This field is in phase with the pole variation. Now the magnetization will produce a sheet of
poles on the lower surface of the magnetized slab, and the field produced by this surface
will be directly opposite to the field produced by the upper surface, since the poles of the
upper and lower surfaces are of opposite sign. Hence the field produced by a slab of finite
thickness will be
2nA sin (ky‘) [exp (- kh) - exp (- k(h + d ) ) l ,
where d is the thickness of the slab. Therefore the power spectrum of the field is given by
P’(k) =
2pA2
p2 + (kZs2/4)
. 4n2 [exp(-kh)
-
exp(-k(h + d ) ) ] ’
- 8n2pA2exp (- 2kh) [ 1 - exp (- kd)]
p2 + (kZs2/4)
The quantity 2n[exp (-kh)-exp ( 4(h + d ) ) ] has also been derived by Schouten &
McCamy (1972), who called it the Earth fdter.
At large values of k, an approximation for the power spectrum is
P‘(k) =
32nZpA2exp (- 2kh)
k2sZ
Therefore log, [P’(k)] = c - 2 log, k - 2kh,
where c is a constant.
2 .
Therefore - log, [PI@)] = - - - 2h
ak
k
’1
1
confirming Spector & Grant’s (1970) derivation for the slope at high wavenumbers.
The possible location of the source of the magnetic anomalies investigated by Blakely
el ul. (1973) was studied using the theoretical power spectrum derived above. The slope of
the spectrum calculated by Blakely et ul. (1973) was -10.3 between wavenumbers of 0.1
and 0.3 rad km-’,or about twice the depth that the oceanic basement was below the
aircraft. This immediately tells us that the average depth to the source of the anomalies must
be greater than the top of the basement, since the form of the theoretical spectrum is for
the negative slope to increase at wavenumbers larger than 0.3 rad km-’ . The shape of the
power spectrum has been analysed as a function of h, the depth to the top of the slab;
d, the thickness of the slab; and p , the reversal rate.
276
C. G. A . Harrison
Firstly, the variation with reversal rate is presented in Fig. 6. The depth to the top surface
is taken as 5 km and the thickness is taken as 0.5 km. The effect of increasing the reversal
rate is to make the negative slope more shallow in the region between 0.1 and 0.3 rad km-'.
In other words, if there are more undetected polarity events within this time period the
slope will become smaller. Assuming that the real reversal rate is that shown by model A of
Blakely et al. (1973, Fig. 5) in which there are 30 reversals in 19 Ma, we can see that if the
source of magnetic anomaly were in fact residing in the top 500 m of the oceanic basement,
we would expect a slope of -7.3, which is smaller in magnitude than the observed slope of
-10.3. Even if model B is taken and reversals are added at A l , A2 and A3, we have a reversal
rate of 1.26 Myr-I, which is not small enough to give the required negative slope of the
observed spectrum.
90-
h= 5 0 km
;80W
3
-z
70-
0
60-
5
W
b
50-
/u
2 21
I58
2
40-
095
30-1
Secondly, the variation of the power spectrum as the depth to the surface is changed is
shown in Fig. 7. This shows clearly that the effect of increasing h is to steepen the negative
slope of the spectrum in the region of interest. Finally, the effect of increasing the thickness
of the slab is shown in Fig. 8. Increasing the thickness of the slab will also cause a steepening
of the slope in the region of interest, although this is not as efficient a method of increasing
the slope as that used in the previous figure.
In summary, we can state that in order to achieve a negative slope of 10.3 for the power
spectrum either one or more of the following things has to occur: (a) The reversal rate has to
be much less than that measured, (b) The depth to the top of the body causing the magnetic
anomalies has to be deeper than the surface of the basement, (c) The thickness of the body
has to be much greater than 0.5 km. We reject the first alternative, not only because of
previous discussion of this matter but also because of the shape of the power spectrum
itself. If Figs 6-8 are studied it will be seen that the peak of the power spectrum is affected
mainly by the reversal rate, moving to larger wavenumbers as the reversal rate is increased.
The peak in the observed power spectrum is at about 0.1 rad k m - I , almost exactly where it is
in the theoretical spectrum for a reversal rate of 1.58 Myr-'. In order to achieve a peak
power at this wavenumber for a smaller reversal rate, the depth would have to be less than
5 km and/or the thickness would have to be less than 0.5 km, neither of which appears to be
in the least likely. Hence we conclude that the depth must be greater or the thickness greater
Magnetization of the oceanic crust
277
9.0
8.0
I
-f
W
O'
6.0
s
z
a
5.0
$
4.0
E
z
30
1
1
1
1
I
I
1
2
3
4
5
6
WAVENUMBER. RADIANS / KM
Figure 7. Curves of the variation of natural logarithm of power as a function of wavenumber for different
depths to the top surface. The thickness of the layer and the reversal rate are kept constant.
12 0
h i 5 0 km
,u= I 5 8 I Ma
I1 0
a
-:
10 0
so
(3
0
-I
80
z
9
K
g ro
a
d
20
z
60
I .o
50
0 5
40
1
2
3
4
5
6
WAVENUMBER. RADIANS / KM
Figure 8. Curves of the variation of natural logarithm of power as a function of wavenumber for different
thicknesses. The depth to the top surface and the reversal rate are kept constant.
than for the commonly accepted model. It is possible to calculate pairs of depth and
thickness values which give the required slope of -10.3 between wavenumbers of 0.1 and
0.3 rad km-'. These are given in Table 6. This result is in agreement with the result obtained
by the study of the intensity of anomalies and the intensity of magnetization in the upper
oceanic crust, in that we have suggested that deeper layers are involved in the formation of
the marine magnetic anomalies.
Rutten (1975) has also computed power spectra for oceanic magnetic anomalies. The
data which he obtained over the Reykjanes Ridge gave a power spectrum in which the slope
of the natural logarithm of the power plotted against wavenumber was about -12.4,
suggesting an average depth to the source of more than 6.2 km, which is much deeper than
278
C. G. A. Ham'son
Table6. Depth ( h ) and thickness (d)values necessary to give a slope of
using a reversal rate of 1.58 reversalsMa-'
d (km)
h (km)
0.5
6.48
1.0
6.24
2.0
5.79
- 10.3
between 0.1 and 0.3 rad km-I
3.0
5.38
Note:
These results show that the depth to the middle of the layer (h + d/2)remains approximately constant
at about 6.8 km.
the depth suggested by the calculations of Talwani et al. (1971). These calculations by
Rutten confirm the calculations presented here, concerning the depth of the source.
Speculations as to the deeper source
The extrusive and moderately highly-magnetized layer is unlikely to be greater than 1 km
thick. It is difficult to see how larger thicknesses of extrusives could be built up in a region
which is undergoing continuous spreading. Also, it seems likely that hydrothermal metamorphism which is believed to occur in the oceanic crust would metamorphose lower
basaltic layers, with the concomitant severe reduction in the intensity of magnetization (Fox
& Opdyke 1973). Also the seismic evidence for the thickness of layer 2A, which has been
equated with the extrusive layer, suggests an average thickness of less than 1km.
Other rock types exist which could be important for producing magnetic anomalies. Our
knowledge of the magnetic properties of these other rocks is based almost entirely on
dredged samples, which we have suggested do not give a representative picture in the case of
the basalts. However, since these are the only samples available, we shall discuss the results
from them, with the proviso that they may not give an adequate representation of the lower
layers of the oceanic crust. Gabbro is often considered to be an important constituent of the
oceanic crust, especially layer 3 (e.g. Christensen & Salisbury 1975). Data presented by Fox
& Opdyke (1973) suggests that gabbros have stable magnetizations, and that if they are not
weathered, their average magnetization is in the region of 1 Am-'. Thus if fairly large
percentages of the oceanic crust were formed of unweathered gabbro, this rock type could
contribute significantly to marine magnetic anomalies. However, this value for the magnetization of gabbro may be too high. Carmichael (1970) has studied six coarse-grained rocks
from the mid-Atlantic ridge and finds intensities of magnetization which are one half or less
than the average value given by Fox & Opdyke (1973). The one exception in Carmichael's
set of samples has a very high magnetization, but he thought that this sample was a highlyweathered surface sample, and not originally coarse-grained. In addition, Irving et al. (1970)
give a mean value for three gabbro and diabase samples of only one tenth that of Fox &
Opdyke. However, the gabbros of Fox & Opdyke are quite stable towards alternating field
demagnetization, and so could be carriers of stable magnetization. The one gabbro studied
by Opdyke & Hekinian (1967) had a magnetization of only 0.5 A m-' but was very stably
magnetized.
Another type of rock which is a popular contender in many models is amphibolite. Very
few samples of amphibolite have had their magnetic properties studied. Irving et al. (1970)
studied four samples of amphibolite and found magnetization intensities between 1.3 A m-l
and 1 mA m-'. The arithmetic mean value was 0.63 A m-', but the scatter is so great that we
essentially cannot make any predictions as to the possible importance of amphibolite on
marine magnetic anomalies.
Magnetization of the oceanic crust
279
A third possibility is that serpentinized peridotite is an important constituent of the
oceanic crust. Oceanic serpentinites are known to have very high values of magnetization
sometimes. Fox & Opdyke (1973) list three serpentinites with an arithmetic mean NRM of
8.2 Am-'. Irving et al. (1970) analysed five samples of serpentinite. The largest value of
NRM was 7.5 A m-' and the arithmetic mean was 3.2 A m-'. The serpentinites of Fox &
Opdyke had median demagnetizing inductions of 0.14 mT, and larger, which are high
enough values to suggest reasonable stability of magnetization.
A great deal more is known about continental serpentinites. Saad (1969) studied ultramafic rocks from the Franciscan formation and found that both susceptibility and NRM
intensity increased with decreasing density. The decrease of density is correlated with
increasing degree of serpentinization. The minimum density of lus samples was about
2.55 x 103kg m-3, and at these densities, the intensity of magnetization was 1.5 A m-'.
He also found out that the highly-serpentinized samples were less stably magnetized than the
other samples. Hatherton (1 967) also found a negative correlation between susceptibility
and density for ultramafics, but did not measure remanent magnetization. Cox, Doell &
Thompson (1964) found a positive correlation between density and both susceptibility and
remanent magnetization in the serpentinite recovered from the AMSOC hole in Puerto Rico.
This behaviour appears to be anomalous, as was suggested by Saad (1969). It may be associated with the removal of magnetite by secondary weathering phenomena, after the olivine
has been completely serpentinized (Watkins & Paster 1970). Komarov et al. (1962) also
show increasing intensity of magnetization with increasing degree of serpentinization for
ultramafics from the Urals. The maximum intensity of magnetization was 17.5 A m-', but
no information was given concerning the stability of magnetization.
The amount of serpentinite in the oceanic crust is a matter of some controversy. It was
of course a favourite constituent of Hess (1962). Serpentinite is part of Steinmann's (1906)
Trinity of Alpine ophiolites, which are believed to represent sections of the oceanic crust.
Bonatti & Honnorez (1976) believe that serpentinite is a significant constituent of layer 3.
In fracture zones of the equatorial Atlantic, serpentinites are one of the mzjor rock types
dredged from the scarps. Some of the scarps are obviously anomalous, but Bonatti &
Honnorez make the point that the northern wall of the Vema fracture zone appears to have
the characteristics of normal oceanic crust, and on this wall serpentinites are commonly
dredged. Christensen (1972) has analysed secondary wave velocities (V,) observed in the
oceanic crust. In some marine seismic refraction experiments, what are believed to be
converted shear waves are observed occasionally as second arrivals. If this interpretation is
correct, then it is possible to measure the V, velocity of layer 3 of the oceanic crust.
Knowledge of Vp from the Same experiment then allows the Poisson's ratio of layer 3 to be
calculated. Values of the Poisson's ratio obtained in this way vary between 0.21 and 0.29.
A Poisson's ratio of 0.29 gives a V,/V, ratio of 1.84. Christensen's (1972) measurements
of seismic velocities of partially and completely serpentinized peridotites show that if a
partially serpentinized peridotite is chosen to give the correct V,, of 6.7 km s-' for layer 3
of the oceanic crust, then the V,,/V, ratio will be 1.91. In other words, the shear-wave
velocity of serpentinites with the correct pressure-wave velocity for layer 3 are much lower
than what is observed for the oceanic crust.
These results of Christensen, however, do not mean that there is no serpentinite in
layer 3. Seismic refraction measurements produce average values of velocities and the
division into layers of constant seismic velocity is to some extent an artifact of the method
of analysis. It does not mean that these layers are composed of homogeneous rocks.
Evidence of direct sampling suggests that each layer is composed of conglomerations of
many different rock types. The presence of up to 20 per cent serpentinite in the oceanic
280
C.G. A . Ham'son
layer 3 might go undetected seismically, although it would be extremely important
magnetically.
The model of Cann (1974) requires a magma chamber situated below the ridge crest
about 1.5 km below the surface, and about 4.5 km thick. This magma chamber produces
extrusive lava at the surface, by way of feeder dykes rising from the upper portion of the
magma chamber. This lava-dyke complex gives rise to seismic layer 2. Cann believes that
cumulate layers form at the bottom of the magma chamber by settling of crystals which
solidify within the magma chamber, in much the same way as cumulate layers are formed at
the base of layered continental intrusives such as the Stillwater, Bushveldt and Skaegaard
intrusive complexes. He also believes that gabbroic rocks crystallize out directly on the walls
of the magma chamber due to the cooling of the chamber with time. It is the gabbroic rocks
and some of the layered complexes which form seismic layer 3. However, if the magma
chamber is of basaltic composition, as it must be to produce extrusive tholeiitic basalts, then
any crystal which is in equilibrium with the liquid will have to be more ultrabasic than the
magma. Hence it seems perfectly possible that ultramafic rocks could crystallize on the sides
of the magma chamber.
A serpentinite model for the production of magnetic anomalies does encounter severe
difficulties. The first is that the olivine has to be serpentinized, which needs the presence of
water. It is difficult to know exactly where this serpentinization occurs. There is a growing
body of evidence that suggests significant circulation of sea water through the upper
portions of the oceanic crust, but whether this circulation penetrates into layer 3 is not
known. If, however, enough water is present, then the time of serpentinization will be
governed by the time at which the peridotite +water system cools below 5OO0C, this being
the temperature above which serpentinite becomes dehydrated. As pointed out by Cande &
Kent (1976) this isotherm slopes downward away from the axial zone, so that the zones
of normally- and reversely-magnetized serpentinite will be separated by sloping interfaces.
The second problem is that for serpentinite to be an important source for sea floor
spreading anomalies, it has to occur more or less uniformly either as a layer in the manner
suggested by Cann (1974) or dispersed throughout layer 2 or 3. However, many people
believe that serpentinite is tectonically emplaced within the oceanic crust, along faulted
boundaries. If this is the case, then we would not expect to find continuous enough material
to produce the very regular sea floor spreading anomalies which are observed. It also appears
possible that tectonic emplacement of serpentinites will disturb the original magnetization
sufficiently so any anomaly observed over such a body would have little relationship to the
original direction of magnetization.
Conclusions
The intensities of magnetization in basalts recovered by the DSDP are small. In order to
produce the sea floor spreading anomalies from such rocks, a thickness of several kilometres
is necessary.
Independent evidence for a magnetized layer several kilometres thick can be obtained
from the power spectrum of marine magnetic anomalies. Since extrusive basalts are probably
limited to layers less than 1 km thick, other rock types must be important in the formation
of marine magnetic anomalies. The magnetic properties (as far as they are known) of the
various rock types which might make up the oceanic crust are discussed. Unfortunately, too
little is known about the magnetic properties of these rocks to enable us to decide which one
is mainly responsible.
Magnetization of the oceanic crust
28 1
Acknowledgments
I have benefited from discussions with J. Honnorez and E. Bonatti. Research was supported
by the National Science Foundation.
Contribution from the University of Miami, Rosenstiel School of Marine and Atmospheric
Science.
References
Atwater, T. & Mudie, J. D., 1973. Detailed near-bottom geophysical study of the Gorda rise, J. geophys.
Res., 78, 8665-8686.
Blakely, R. J. & Cox, A., 1972. Evidence for short geomagnetic polarity intervals in the Early Cenozoic,
J. geophys. Res.. 77,7065-7072.
Blakely, R. J., Cox, A. & Eufer, E. J., 1973. Vector magnetic data for detecting short polarity intervals
in marine magnetic profdes, J. geophys. Res., 78,6977-6983.
Bonatti, E. & Honnorez, J., 1976. Sections of the Earth's Crust in the Equatorial Atlantic, J. geophys.
Res.. in press.
Bott, M. H. P., 1967. Solution of the linear inverse problem in magnetic interpretation with application
to oceanic magnetic anomalies, Geophys. J. R. astr. SOC.,13,313-323.
Bott, M. H. P. & Hutton, M. A., 1970. A matrix method for interpreting oceanic magnetic anomalies,
Geophys. J. R . astr. Soc., 20, 149-157.
Butler, R. F. & Banerjee, S. K., 1973. Magnetic properties of exposed oceanic crust on Macquarie
Island, Nature Phys. Sci., 244,115-118.
Butler, R. F., Banerjee, S. K. & Stout, J. H., 1975. Magnetic properties of oceanic pillow basalts from
Macquarie Island, Nature, 257,302-303.
Cande, S. & Kent, D. V., 1976. Constraints imposed by the shape of marine magnetic anomalies on the
magnetic source, J. geophys. Res., in press.
Cann, J. R., 1974. A model for oceanic crustal structure developed, Geophys. J. R. astr. SOC.,39,
169-1 87.
Carmichael, C. M., 1970. The mid-Atlantic ridge near 4 5 " N . VII. Magnetic properties and opaque
mineralogy of dredged samples, Can. J. Earth Sci., 7,239-256.
Christensen, N. I., 1972. The abundance of serpentinites in the oceanic crust. J. Geol., 80,709-719.
Christensen, N. I. & Salisbury, M. H., 1975. Structure and constitution of the lower oceanic crust, Rev.
Geophys. Space Phys., 13,57-86.
Cox, A. & Doell, R. R., 1962. Magnetic properties of the basalt in hole EM7, Mohole Project, J. geophys.
Res., 67,3997-4004.
Cox, A., Doell, R. R. & Thompson, G. A., 1964. Magnetic properties of serpentinite from Mayaguez,
Puerto Rico, A study of serpentinite, Nat. Acad. Sci., Nat. Res. Council Publ., 1188.
Emilia, D. A. & Bodvarsson, G., 1969. Numerical methods in the direct interpretation of marine magnetic
anomalies, Earth Planet. Sci. Lett., 7, 194-200.
Fox, P. J. & Opdyke, N. D., 1973. Geology of the oceanic crust: magnetic properties of oceanic rocks,
J. geophys. Res.. 78,5 139-5 154.
Francheteau, J., Harrison, C. G. A., Sclater, J. G . & Richards, M. L., 1970. Magnetization of Pacific Seamounts: a preliminary polar curve for the Northeastern Pacific, J. geophys. Res., 75,2035-2061.
Harrison, C. G. A., 1968. Formation of magnetic anomaly patterns by dyke injection, J. geophys. Res.,
73,2137-2142.
Harrison, C. G. A., 1970. Magnetization of Atlantic seamounts, Bull. mar. Sci., 20,560-574.
Harrison,C. G. A., 1971. A seamount with a nonmagnetic top, Geophys., 36,349-357.
Harrison, C . G. A. & Ball, M. M., 1974. Geophysical observations on an exposed seamount in the Afar
depression, Bull. Volcano., 38,26-43.
Harrison, C. G. A., Jarrard, R., Vacquier, V. & Larson, R. L., 1975. Paleomagnetism of Cretaceous Pacific
Seamounts, Ceophys. J. R . astr. SOC..42,859-882.
Hatherton, T., 1967. A geophysical study of Nelson-Cook Strait region, New Zealand, J. Geol. Geophys.
N.Z., 10,1330-1347.
Heirtzler, J. R., Dickson, G . O., Herron, E. M., Pitman, W. C. & Le Pichon, X., 1968. Marine magnetic
anomalies, geomagnetic field reversals, and motions of the ocean floor and continents, J. geophys.
Res., 73,2119-2136.
C. G. A. Harrison
282
Herron, E. M., 1972. Sea-floor spreading and the Cenozoic history of the EastCentral Pacific, Geol. SOC.
Am. Bull., 83,1671-1692.
Hess, H. H., 1962. History of ocean basins, Petrologic studies: a volume to honor A. F. Buddington, eds
A. E. J. Engel, H. L. James & B. F. Leonard, pp. 599-620. Geol. SOC.Am.
Hyndman, R. D., Aumento, F., Melson, W. G., Hall, J. M., Bougault, H., Dmitriev, L., Fisher, J. F.,
Flower, M., Howe, R. C., Miles, G. A., Robinson, P. T. & Wright, T. L., 1976.Seismic structure of
the oceanic crust from deep drilling on the mid-Atlantic ridge, Geophys. Res. Lett., 3,201-204.
Irving, E., Robertson, W. A. & Aumento, F., 1970. The mid-Atlantic &dge near 45" N. VI. Remanent
intensity, susceptibility, and iron content of dredged samples, Con. J. Earth Sci., 7,226-238.
Klitgord, K. D., 1974. Near-bottom geophysical surveys and their implications on the crustal generation
process, sea-floor spreading history of the Pacific, and the geomagnetic time scale: 0 to 6
M.Y.B.P., unpublished PhD dissertation, University of California, San Diego.
Komarov, A. G., Moskaleva, S. V., Belyayev, V. M. & Uyina, V. I., 1962.Interpretation of magnetic fields
over ultrabasic complexes: serpentinization and magnetic properties, Doklady Akad. Nauk SSSR,
143,70-71.
Larson, R. L., Larson, P. A., Mudie, J. D. & Spiess, F. N., 1974. Models of near-bottom magnetic
anomalies on the East Pacific Rise crest at 21" N, J. geophys. Res., 79,2686-2689.
Lowrie, W., 1974.Oceanic basalt magnetic properties and the Vine and Matthews hypothesis, J. Geophys.,
40,513-516.
Lowrie, W., h v l i e, R. & Opdyke, N. D., 1973a. The magnetic properties of deep sea drilling project
basalts from the Atlantic Ocean, Earth Planet Sci. Lett., 17,338-349.
Lowrie, W., b v l i e , R. & Opdyke, N. D., 1973b. Magnetic properties of Deep Sea Drilling Project Basalts
from the North Pacific Ocean, J. geophys. Res., 78,7647-7660.
Luyendyk, B. P., 1969. Origin of short-wavelength magnetic lineations observed near the ocean bottom,
J. geophys. Res., 74,4869-4881.
McEJhinny, M. W., 1973.Palaeomagnetism and plate tectonics, 358 pp. Cambridge University Press.
McKenzie, D. & Sclater, J. G., 1971. The evolution of the Indian Ocean since the Late Cretaceous,
Geophys. J. R. astr. SOC.,24,437-528.
Matthews, D. H., 1961. Lavas from an abyssal hill on the floor of the N. Atlantic Ocean, Nature, i90,
158-159.
Matthews, D. H. & Bath, J., 1967. Formation of magnetic anomaly pattern of mid-Atlantic ridge,
Geophys. J. R . astr. Soc., 13,349-357.
Opdyke, N. D. & Hekinian, R., 1967. Magnetic properties of some igneous rocks from the mid-Atlantic
ridge, J. geophys. Res.. 72,2257-2260.
Parker, R. L. & Heustis, S. P., 1974. The inversion of magnetic anomalies in the presence of topography,
J. geophys. Res.. 79,1587-1593.
Peter, G., 1970. Discussion of paper by B. P. Luyendyk, Origin of short-wavelength magnetic lineations
observed near the ocean bottom, J. geophys. Res., 75,6717-6720.
Pitman, W. C., Larson, R. L. & Herron, E. M., 1974.Age of the ocean basins determined from magnetic
anomaly lineations, Two maps, published by the Geological Society of America.
Pitman, W. C. & Talwani, M., 1972. Sea floor spreading in the North Atlantic, Geol. SOC.Am. Bull., 83,
619-646.
Poehls, K. A., 1974. Seismic Refraction on the mid-Atlantic Ridge at 37"N, J. geophys. Res., 79,
3370-3373.
Rice, S . O., 1954. Mathematical theory of random noise, Selected papers on noise and stochastic processes, pp. 133-294, ed. Nelson Wax, Dover Publications Inc., New York.
Richards, M. L., Vacquier, V. & Van Voorhis, G. D., 1967. Calculation of the magnetization of uplifts
from combining topographic and magnetic surveys, Geophys., 30,797-817.
Rutten, K., 1975. Twodimensionality of magnetic anomalies over Iceland and Reykjanes ridge, Mar.
geophys. Res., 2,243-263.
Saad, A. H., 1969. Magnetic properties of ultramafic rocks from Red Mountain, California, Geophys.,
34,974-987.
Schouten, H. & McCamy, K., 1972. Filtering marine magnetic anomalies, J. geophys. Res., 77,
7089-7099.
Scientific Party Deep Sea Drilling Project Leg 37, 1975. Sources of magnetic anomalies on the midAtlantic Ridge, Nature, 255,389-390.
Spector, A. & Grant, F. S., 1970. Statistical models for interpreting aeromagnetic data, Geophys., 35,
293-302.
Magnetization of the oceanic crust
283
Steinmann, G., 1906. Geologische Beobachtungen in den Alpen (11). Die Schartsche ijberfaltungstheorie
und die geologische Bedeutung der Tiefseeabsatze und der ophiolithischen Massengesteine, Ber.
Naturh Ges. Freiburg i. Br., 16, 1-49.
Talwani, M., Windisch, C. C. & Langseth, M. G., 1971. Reykjanes ridge crest: a detailed geophysical
study, J. geophys. Res., 76,473-517.
Tarasiewicz, E., Tarasiewicz, G. & Harrison, C. G. A., 1976. Magnetic properties of leg 37 basalts, Initiul
Reports of the Deep Sea Drilling Project, leg 37, eds R. S. Yeats & S. R. Hart, 473-478.
Van den Akker, F. B., Harrison, C. G. A. & Mudie, J. D., 1970. Even more on the direct interpretation
of magnetic anomalies, Earth Planet. Sci. Lett., 9,405 -406.
Vine, F. J., 1968. Paleomagnetic evidence for the northward movement of the North Pacific basin during
the past 100 m.y., Trans. Am. geophys. Un., 49,156.
Vine, F. J. & Hess, H. H., 1970. Sea floor spreading, The Sea, Vol. IV part 2, pp. 587-622, ed. A. E.
Maxwell, Wiley-Interscience.
Watkins, N. D., 1972. Review of the development of the geomagnetic polarity time scale and discussion
a f prospects for its finer definition, Geol. SOC.Am. Bull., 83,551-574.
Watkins, N. D. & Paster, T. P., 1970. The magnetic properties of igneous rocks from the ocean floor,
Phil. Trans. R. SOC.Lond. A , , 268,507-550.