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Note. The second plate moves clockwise relative to the first plate.

Note. The second plate moves clockwise relative to the first plate.

when (p = 90°. For most ridges (e.g., the Cocos-Pacific ridge) cp never approaches 90° and the fastest actual spreading rate today is about 150 km Myr"1 along the East Pacific Rise separating the Nazca and Pacific plates.

5.1.2 Vine-Matthews Crustal Model

Newly erupted ocean-floor spreads smoothly away from the mid-ocean ridges. Because these submarine basalts cool quickly from high temperatures in the sea water environment, they contain fine-grained titanomagnetite of average composition Fe2.4Ti06O4 (TM60). This quenching process prevents high-temperature deuteric oxidation, as generally occurs in subaerial basalts, so that typical Curie temperatures lie in the range 150-200°C. Rapid cooling is also responsible for the fine grain size, enabling oceanic basalts to acquire an intense TRM and become excellent recorders of the paleomagnetic field.

Vine and Matthews (1963) and Morley and Larochelle (1964) proposed that the lineated magnetic anomalies (magnetic stripes), which are observed parallel to and on either side of the mid-ocean ridge, are records of past reversals of the geomagnetic field. They therefore modeled the magnetic anomalies as being produced by strips of crust of alternating polarity at increasing distances (and therefore age) from the ridge crest (Fig. 5.5). For a uniform spreading rate the strips, and their linear magnetic anomalies at the ocean surface, will have widths related to the durations of the geomagnetic polarity epochs. In analyzing these anomalies it is usual to use the effective susceptibility of the material according

Gilbert' Gauss

Gauss

Matuyama

Brunhcs

Matuyama

Fig. 5.5. Schematic representation of sea-floor spreading and the formation of linear magnetic anomalies due to reversals of the Earth's magnetic field as proposed by Vine and Matthews (1963). Normal polarity zones are shaded. Only major subchrons are shown. Updated after Allan (1969), with the permission of Elsevier Science.

Gilbert' Gauss

Gauss

Gilbert

Matuyama

Brunhcs

Matuyama

Fig. 5.5. Schematic representation of sea-floor spreading and the formation of linear magnetic anomalies due to reversals of the Earth's magnetic field as proposed by Vine and Matthews (1963). Normal polarity zones are shaded. Only major subchrons are shown. Updated after Allan (1969), with the permission of Elsevier Science.

to (2.1.3) given by the total magnetization (remanent plus induced) as in (2.1.2) divided by the present magnetic field strength. However, the magnetic properties of rock samples dredged from the ocean floor demonstrated the predominance of remanent magnetization since the actual susceptibility is comparatively low. For basalts the effective susceptibility is of the order of 0.1 SI units, corresponding to a magnetization of about 5 Am"1.

5.1.3 Measurement of Marine Magnetic Anomalies

Marine magnetic anomalies are generally measured by towing a total-field magnetometer on the ocean surface behind a ship at sufficient distance so that the magnetic effects of the ship are below the noise level of the magnetometer. Anomalies may also be measured by flying an aircraft equipped with a suitable magnetometer at a known height above the ocean surface. Satellite measurements may also be used but the filtering effect of the height above the ocean floor makes the results less valuable because most of the detail is lost. Most ship-borne magnetometers are proton precession or fluxgate types that only measure the total intensity F of the geomagnetic field. At any point on the

Earth's surface the expected value of the total intensity (F0) may be calculated using the International Geomagnetic Reference Field (IGRF) for the appropriate epoch, as defined in §1.1.3. The magnetic anomaly (AF) is then given by

It is worth noting that this does not necessarily give exactly the same value for the anomaly as would be obtained using vector observations and calculations.

Ocean-surface magnetometers measure the geomagnetic field up to 5 km or more above the ocean floor, but sometimes deep-tow magnetometers, which follow the bottom topography of the ocean floor at a fixed distance above it (generally in the range 50-200 m), are used (Luyendyk et al., 1968; Greenewalt and Taylor, 1974; Macdonald, 1977; Sager et al., 1998). Parker (1997) discussed the analysis and interpretation of measurements from two deep-tow magnetometers in which the second magnetometer is towed 300 m above the first, giving in effect a vertical gradiometer. Deep-tow magnetometers provide much more detail that those that measure at the ocean surface because they are nearer to the signal source. An example of this difference is shown in Fig. 5.6, in which the detail measured by the deep-tow near the ocean floor is lost when measured at the surface. In the early days of interpreting marine magnetic

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