Gravity Anomaly Over Oceanic Ridge

Fig. 6.3 (a) Axial relief and (b) seismic crustal thickness as a function of full spreading rate at mid-ocean ridge crests. A ridge classification scheme is shown by the heavy black straight lines which indicate the spreading rate ranges for ultraslow, slow, fast and two intermediate classes (modified from Dick et al., 2003, with permission from Nature 426, 405-12. Copyright © 2003 Macmillan Publishers Ltd).

20 40 60 80 100 120 Full spreading rate (mm yr-1)

Fig. 6.3 (a) Axial relief and (b) seismic crustal thickness as a function of full spreading rate at mid-ocean ridge crests. A ridge classification scheme is shown by the heavy black straight lines which indicate the spreading rate ranges for ultraslow, slow, fast and two intermediate classes (modified from Dick et al., 2003, with permission from Nature 426, 405-12. Copyright © 2003 Macmillan Publishers Ltd).

6.2 BROAD STRUCTURE OF THE UPPER MANTLE BELOW RIDGES

Gravity measurements have shown that free air anomalies are broadly zero over ridges (Figs 6.4, 6.5), indicating that they are in a state of isostatic equilibrium (Section 2.11.6), although small-scale topographic features are uncompensated and cause positive and negative free air anomalies. The small, long wavelength, positive and negative free air anomalies over the crests and flanks, respectively, of ridges are a consequence of the compensation, with the positives being caused by the greater elevation of the ridge and the negatives from the compensating mass deficiency. The gravitational effects of the compensation dominate the gravity field away from the ridge crest, and indicate that the compensation is deep.

Seismic refraction experiments by Talwani et al. (1965) over the East Pacific Rise showed that the crust is slightly thinner than encountered in the main ocean basins, and that the upper mantle velocity beneath the crestal region is anomalously low (Fig. 6.4). Oceanic layer 1 rocks (Section 2.4.5) are only present within topographic depressions, but layers 2 and 3 appear to be continuous across the ridge except for a narrow region at the crest. A similar structure has been determined for the Mid-Atlantic Ridge (Fig. 6.5). The suggestion of this latter work that layer 3 is not continuous across the ridge was subsequently disproved (Whitmarsh, 1975; Fowler, 1976).

As the crust does not thicken beneath ridges, iso-static compensation must occur within the upper mantle by a Pratt-type mechanism (Section 2.11.3). Talwani et al. (1965) proposed that the anomalously low upper mantle velocities detected beneath ridges correspond to the tops of regions of low density. The densities were determined by making use of the Nafe-Drake relationship between P wave velocity and density (Nafe & Drake, 1963), and a series of models produced that satisfied both the seismic and gravity data. One of these is shown in Fig. 6.6, and indicates the presence beneath the ridge of a body with a density contrast of -0.25 Mg m-3 extending to a depth of some 30 km. This large density contrast is difficult to explain geologically. An alternative interpretation, constructed by Keen & Tramontini (1970), is shown in Fig. 6.7. A much lower,

Fig. 6.4 Heat flow, free air gravity anomaly and crustal structure defined by seismic refraction across the East Pacific Rise at 15-17°S. P wave velocities in km s~' (redrawn from Talwani et al., 1965, by permission of the American Geophysical Union. Copyright © 1965 American Geophysical Union).
Gravity Anomaly Along The Ridge

500 0 500 1000 km

Fig. 6.5 Gravity anomalies and crustal structure defined by seismic refraction across the Mid-Atlantic Ridge at about 31°N. Bouguer anomaly reduction density 2.60Mg m~3, P wave velocities in km s- (redrawn from Talwani et al., 1965, by permission of the American Geophysical Union. Copyright © 1965 American Geophysical Union).

500 0 500 1000 km

Fig. 6.5 Gravity anomalies and crustal structure defined by seismic refraction across the Mid-Atlantic Ridge at about 31°N. Bouguer anomaly reduction density 2.60Mg m~3, P wave velocities in km s- (redrawn from Talwani et al., 1965, by permission of the American Geophysical Union. Copyright © 1965 American Geophysical Union).

Gravity Anomalies Across Mor
Fig. 6.6 Possible model of the structure beneath the Mid-Atlantic Ridge from gravity modeling with seismic refraction control. Densities in Mg m- (redrawn from Talwani et al., 1965, by permission of the American Geophysical Union. Copyright © 1965 American Geophysical Union).

more realistic density contrast of -0.04 Mg m-3 is employed, and the anomalous body is considerably larger, extending to a depth of 200 km. However, this model can also be criticized in that the densities employed are rather too high, and provide too low a density contrast, and the depth to the base of the anomalous mass is too great. A model that employs densities of 3.35 and 3.28 Mg m-3 for normal and anomalous mantle, respectively, with the anomalous mass extending to a depth of 100 km, would be more in accord with geologic and geophysical data. Indeed, seismic tomography (Section 2.1.8) suggests that the low velocity region beneath ocean ridges extends to a depth of 100 km (Anderson et al., 1992).

Given the ambiguity inherent in gravity modeling, the two interpretations shown probably represent end

Gravity Anomaly Along The Ridge

500 0 500 km

Fig. 6.7 Alternative model of the structure beneath the Mid-Atlantic Ridge from gravity modeling. Profile at 46°N. Densities in Mgm- (redrawn from Keen & Tramontini, 1970, with permission from Blackwell Publishing).

500 0 500 km

Fig. 6.7 Alternative model of the structure beneath the Mid-Atlantic Ridge from gravity modeling. Profile at 46°N. Densities in Mgm- (redrawn from Keen & Tramontini, 1970, with permission from Blackwell Publishing).

members of a suite of possible interpretations. They demonstrate without ambiguity, however, that ridges are underlain by large, low-density bodies in the upper mantle whose upper surfaces slope away from the ridge crests.

6.3 ORIGIN OF ANOMALOUS UPPER MANTLE BENEATH RIDGES

There are three possible sources of the low-density regions which underlie ocean ridges and support them isostatically (Bott, 1982): (i) thermal expansion of upper mantle material beneath the ridge crests, followed by contraction as sea floor spreading carries it laterally away from the source of heat, (ii) the presence of molten material within the anomalous mantle,

(iii) a temperature-dependent phase change. The high temperatures beneath ocean ridge crests might cause a transition to a mineralogy of lower density.

Suppose the average temperature to a depth of 100 km below the Moho is 500°C greater at the ridge crest than beneath the flanking regions, the average density to this depth is 3.3 Mg m-3 and the volume coefficient of thermal expansion is 3 X 10-5 per degree. In this case the average mantle density to a depth of 100 km would be 0.05 Mg m-3 less than that of the flanking ocean basins. If isostatic equilibrium were attained, this low-density region would support a ridge elevated 2.2 km above the flanking areas. If the degree of partial melting were 1%, the consequent decrease in density would be about 0.006 Mg m-3. Extended over a depth range of 100 km this density contrast would support a relative ridge elevation of 0.25 km. The aluminous minerals within the upper mantle that might transform to a lower density phase are also the minerals that enter the melt that forms beneath the ridge crest. They are absent therefore in the bulk of the mantle volume under consideration, which consists of depleted mantle; mantle from which the lowest melting point fraction has been removed. It is unlikely then that a phase change contributes significantly to the uplift.

Partial melting of the upper mantle clearly is a reality because of the magmatic activity at ridge crests, but its extent was a matter of conjecture. However, in the mid-1990s a very large-scale experiment, the Mantle Electromagnetic and Tomography (MELT) experiment, was carried out on the crest of the East Pacific Rise specifically to define the vertical and lateral extent of the region of partial melting beneath it (MELT seismic team, 1998). Fifty-one ocean bottom seismometers and 47 instruments that measure changes in the Earth's magnetic and electric fields were deployed across the ridge, between 15° and 18°S, in two linear arrays each approximately 800 km long. This location was chosen because it is in the middle of a long, straight section of the ridge between the Nazca and Pacific plates, and has one of the fastest spreading rates: 146 mm a-1 at 17°S. The extent of any partial melt in the mantle should therefore be well developed in terms of low seismic velocities and high electrical conductivity. Seismic waves from regional and teleseismic earthquakes, and variations in the Earth's electric and magnetic fields, were recorded for a period of approximately 6 months. Analysis of the data revealed an asymmetric region of low seismic velocities extending to a depth of 100 km, with its shallowest point beneath the ridge crest, but extending to 350 km to the west and 150 km to the east of the ridge crest (Fig. 6.8). Both the velocity anomalies and electrical conductivity are consistent with 1-2% partial melting (Evans et al., 1999). There is an indication of incipient melting to a depth of 180 km. The asymmetry of the region of partial melting is thought to be due to a combination of two effects. Within the hot spot framework the western flank of the ridge is moving at more than twice the rate of the eastern flank (Fig. 6.8). It is also close to the South Pacific superswell (Section 12.8.3). Enhanced upwelling and hence flow in the asthenosphere from the superswell and viscous drag beneath the fast moving Pacific plate are thought to produce higher rates of flow and hence higher temperatures beneath the western flank of the ridge. These elevated temperatures are reflected in shallower bathymetry (Section 6.4) and a higher density of sea-mount volcanism on the western flank compared to the eastern flank.

The width of the region of partial melt defined by the MELT experiment seems to be quite wide. One must recall however that the spreading rate at this point is very high, five times higher than that on much of the Mid-Atlantic Ridge. In fact the region of primary melt only underlies crust 2-3 Ma in age, whereas the anomalous uplift of ridges extends out to crust of 70-80 Ma in age. Partial melt in the upper mantle may therefore account for some of the uplift of ridge crests but cannot account for the uplift of ridge flanks.

6.4 DEPTH-AGE RELATIONSHIP OF OCEANIC LITHOSPHERE

The major factor contributing to the uplift of mid-ocean ridges is the expansion and contraction of the material of the upper mantle. As newly formed oceanic lithosphere moves away from a mid-ocean

Distance from axis (km)

West

East

101 mm yr 1

i 1 r Crust

101 mm yr 1

i 1 r Crust

Fig. 6.8 Schematic cross-section beneath the East Pacific Rise at 17°S illustrating the extent of partial melting in the mantle deduced from the results of the MELT experiment. Plate velocities are in the hot spot reference frame. The region labeled E (embedded heterogeneity) indicates enhanced melting due to anomalously enriched mantle or localized upwelling (modified from MELT seismic team, 1998, Science 280,1215-18, with permission from the AAAS).

Fig. 6.8 Schematic cross-section beneath the East Pacific Rise at 17°S illustrating the extent of partial melting in the mantle deduced from the results of the MELT experiment. Plate velocities are in the hot spot reference frame. The region labeled E (embedded heterogeneity) indicates enhanced melting due to anomalously enriched mantle or localized upwelling (modified from MELT seismic team, 1998, Science 280,1215-18, with permission from the AAAS).

ridge, it becomes removed from underlying heat sources and cools. This cooling has two effects. First, the lithosphere contracts and increases in density. Second, because the lithosphere-asthenosphere boundary is controlled by temperature (Section 2.12), the cooling causes the lithosphere to increase in thickness away from the mid-ocean ridge. This latter phenomenon has been confirmed by lithosphere thickness estimates derived from surface wave dispersion studies in the Pacific Ocean, which indicate that the thickness increases from only a few kilometers at the ridge crest to 30 km at 5 Ma age and 100 km at 50 Ma (Forsyth, 1977).

The cooling and contraction of the lithosphere cause a progressive increase in the depth to the top of the lithosphere away from the ridge (Sclater & Francheteau, 1970), accompanied by a decrease in heat flow. It follows that the width of a ridge depends upon the spreading rate, and so provides an explanation for the relative widths of the rapidly spreading East Pacific Rise and more slowly spreading Mid-Atlantic Ridge. Parsons & Sclater (1977) determined the nature of the age-depth relationships of oceanic lithosphere, and suggested that the depth d (meters) is related to age t (Ma) by:

It was found, however, that this relationship only holds for oceanic lithosphere younger than 70 Ma. For older lithosphere the relationship indicates a more gradual increase of depth with age. In order to explain this, Parsons & McKenzie (1978) suggested a model in which the cooling layer comprises two units rather than the single unit implied by Parsons & Sclater (1977). In this model the upper unit, through which heat moves by conduction, is mechanically rigid, and the lower unit is a viscous thermal boundary layer. As the lithosphere travels away from a spreading center, both units thicken and provide the relationship - depth proportional to the square root of age - described above. However, the lower unit eventually thickens to the point at which it becomes unstable and starts to convect. This brings extra heat to the base of the upper layer and prevents it thickening at the same rate. They suggested that the age-depth relationship for oceanic lithosphere older than 70 Ma is then given by:

These two models, for the cooling and contraction of oceanic lithosphere with age, are referred to as the half space and plate models respectively. In the former the lithosphere cools indefinitely, whereas in the latter it ultimately attains an equilibrium situation determined by the temperature at the lithosphere-asthenosphere boundary and the depth at which this occurs as a result of convection in the asthenosphere. Clearly the main constraints on these models are the observed depth (corrected for sediment loading) and heat flux at the ocean floor as a function of age. Stein & Stein (1992), using a large global data set of depth and heat flow measurements, derived a model (GDH1 - global depth and heat flow model 1) that gave the best fit to the observations. Any such model must make assumptions about the depth to the ridge crest and the thermal expansion coefficient, the thermal conductivity, the specific heat, and the density of the lithosphere. However Stein & Stein (1992) showed that the crucial parameters in determining the best fit to the data are the limiting plate thickness and the temperature at the base of the lithospheric plate. In the GDH1 model these have the values 95 km and 1450°C respectively.

A comparison of the age-depth relationship predicted by the half space model, the Parsons, Sclater & McKenzie model and GDH1, is shown in Fig. 6.9a and the depth-age equations for GDH1 are:

d = 2600 + 365t"2 for t < 20 Ma and d = 5650 - 2473exp(-t/36) for t > 20 Ma.

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