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Fig. 6.18 Model for the construction of oceanic crust at a slow-spreading ridge. Transient magma bodies rise to the brittle-ductile transition within the crust and shoulder aside and depress older plutons. Part of the magma body erupts through a fissure to produce a volcano or hummocky lava flow on the sea floor and the remainder solidifies to form part of the main crustal layer (redrawn from Smith & Cann, 1993, with permission from Nature 365,707-15. Copyright © 1993 Macmillan Publishers Ltd).

Fig. 6.18 Model for the construction of oceanic crust at a slow-spreading ridge. Transient magma bodies rise to the brittle-ductile transition within the crust and shoulder aside and depress older plutons. Part of the magma body erupts through a fissure to produce a volcano or hummocky lava flow on the sea floor and the remainder solidifies to form part of the main crustal layer (redrawn from Smith & Cann, 1993, with permission from Nature 365,707-15. Copyright © 1993 Macmillan Publishers Ltd).

overlying, hydrothermally altered dikes into the magma chamber.

The two gabbro units, isotropic and layered, are often correlated with seismic layers 3A and 3B, respectively (Section 2.4.7). The ultramafic cumulates, rich in olivine and pyroxene, would then account for the sub-Moho seismic velocities. Thus, the Moho occurs within the crystallized magma chamber at the base of the mafic section. Off axis, however, in a lower temperature environment, the uppermost ultramafics may become partially hydrated (i.e. serpentinized) and as a result acquire lower seismic velocities more characteristic of layer 3B. The seismic Moho would then occur at a somewhat greater depth, within the ultramafic section. As a result of this uncertainty in defining the seismic Moho, petrologists have tended to define the base of the crust as the base of the presumed magma chamber, that is, the dunite/chromitite horizon. Hence, this level is termed the "petrologic Moho".

The model of Cann (1974) and Kidd (1977) has met with considerable success in explaining the known structure and petrology of oceanic crust created at fast-spreading ridge crests, where there is a steady state magma chamber. At slow-spreading ridge crests, however, the zone of crustal accretion is wider and it seems probable that magma chambers are only tran sient. In this case the alternative model derived from early reinterpretations of ophiolites in terms of sea floor spreading may be more applicable. This invoked multiple small magma chambers within the main crustal layer in the light of the multiple intrusive relationships observed at all levels in the Troodos ophiolite of southern Cyprus (Moores & Vine, 1971) (Section 2.5). Smith & Cann (1993) favor such a model for the creation of oceanic crust at slow-spreading ridge crests (Fig. 6.18). However away from segment centers, and particularly in the vicinity of transform faults, the magma supply may be greatly reduced, and serpentinized mantle peri-dotite appears to be a common constituent of the thinned oceanic crust. This type of crust becomes even more common on very slow-spreading ridges and ultimately most of the crust is effectively exposed mantle with or without a thin carapace of basalts. On the ultraslow Gakkel Ridge the crust is essentially serpentinized and highly tectonized mantle peridotite with volcanic centers at intervals of 100 X 50 km.

An alternative approach to understanding the accre-tionary processes at mid-ocean ridge crests is by way of thermal modeling (Sleep, 1975; Kusznir & Bott, 1976; Chen & Morgan, 1990). Chen & Morgan (1990) made significant improvements to such models by including the effects of hydrothermal circulation at ridge crests and the different rheological properties of the crust compared to the mantle, oceanic crust being more ductile at high temperatures than the mantle. As outlined in Section 6.9, the thermal regime beneath a ridge crest is influenced by the rate at which magma is supplied to the crust, which depends on the spreading rate. As a consequence the brittle-ductile transition (at approximately 750°C) occurs at a shallower depth in the crust at a fast-spreading ridge compared to a slow-spreading ridge that has a lower rate of magma supply. This in turn implies that at a fast-spreading ridge there is a much greater volume, and hence width, of ductile lower crust. This ductile crust effectively decouples the overlying brittle crust from the viscous drag of the con-vecting mantle beneath, and the tensile stresses pulling the plates apart are concentrated in a relatively thin and weak layer that extends by repeated tensile fracture in a very narrow zone at the ridge axis. On a slow-spreading ridge the brittle layer is thicker and the volume of ductile crust much smaller. As a result the tensile stresses are distributed over a larger area and there is more viscous drag on the brittle crust. In this situation the upper brittle layer deforms by steady state attenuation or "necking" in the form of a large number of normal faults creating a median valley.

Chen & Morgan (1990) demonstrated that for crust of normal thickness and appropriate model parameters the transition from smooth topography with a buoyant axial high to a median rift valley is quite abrupt, at a full spreading rate of approximately 70 mm a-1 as observed. The model also predicts that for thicker crust forming at a slow rate of spreading, as for example on the Reyk-janes Ridge immediately south of Iceland, there will be a much larger volume of ductile crust, and smooth topography is developed rather than a rift valley. Conversely, where the crust is thin on a slow-spreading ridge, for example in the vicinity of fracture zones on the Mid-Atlantic Ridge, the median valley will be more pronounced than at a segment center. Such instances of thicker or thinner crust than normal are also likely to be areas of higher or lower than normal upper mantle temperatures respectively which will enhance the effect in each case. The model was extended by Morgan & Chen (1993) to incorporate a magma chamber as observed on the East Pacific Rise. This enhanced model predicts that a steady state magma chamber can only exist at spreading rates greater than 50 mm a-1 and that the depth to the top of the chamber will decrease as spreading rate increases, whilst retaining the essential features of the Chen & Morgan (1990) model.

In general therefore there is good agreement between the theoretical models for the creation of oceanic crust and observations made on in situ ocean floor and on ophiolites. Certain aspects however are still problematic. The evolution of a median valley as accretion occurs, that is, the way in which its flanks are uplifted and the normal faults ultimately reversed, is poorly understood. This is particularly true for the amagmatic segments of very slow- and ultraslow-spreading ridges where mantle material is emplaced directly to the sea floor. The details of the formation of the gabbroic layer 3, from a steady state or transient magma chamber, are also the subject of much debate.

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