The nature of oceanic lower crust and shallow mantle emplaced at low spreading rates

TECTONOPHYSICS ELSEVIER

Tectonophysics 279 (1997) 181-191

The nature of oceanic lower crust and shallow mantle emplaced at low spreading rates N.H. Sleep *, G.A. Barth Department of Geophysics, Stanford University, Stanford, CA 94305-2215, USA Accepted 2 May 1997

Abstract Thermal calculations indicate that at sea floor spreading rates too low to sustain a steady-state magma chamber, most of the upwelling melt still freezes within ~10 km of the sea floor beneath a mid-ocean ridge. A thickened lower crustal section containing crust-mantle mix is thus predicted. Reduced melting in the source region is an expected effect of low spreading rates. However, cooling from the surface limits the degree to which mafic crustal material is separated from mantle at spreading rates higher than those at which reduced melting becomes important. Evidence for ultramafic contributions to the crust at low spreading rates includes a distinct seismic velocity profile for crust from very slow ridges, evidence of aging of the lower crust, spreading-rate-dependent differences in reflection seismic structure and character, and seismic crustal thicknesses apparently greater than adiabatic crustal thicknesses. Crustal structure inferred from the thermal results agrees with the model of Cannat [Cannat, M., 1993. Emplacement of mantle rocks in the seafloor at mid-ocean ridges. J. Geophys. Res. 98, 4163-4172], which includes a lower crust of mixed composition achieved without the requirement of gross tectonic rearrangement of an originally stratified crustal section.

Keywords: mid-ocean ridges; magma chambers; oceanic crust; sea-floor spreading; partial melting; lower crust

1. I n t r o d u c t i o n Geological evidence from sea floor outcrops and drilling results indicates that ultramafic material is sometimes present within the oceanic crustal section (Karson et al., 1987; A u z e n d e et al., 1989; Cannat, 1993; Cannat et al., 1995). This observation is common in crust formed at lower spreading rates, suggesting that the ultramafic component of the crust is volumetrically significant in these settings. There are two contrasting views o f the incorporation of mantle *Corresponding author. Fax: + l norm @pangea.stanford.edu

(415) 725-7344; e-mail:

material into the crustal section at slow-spreading mid-ocean ridges; it m a y be an original component o f the crustal section, or it m a y be e m p l a c e d via gross tectonic rearrangement of normally stratified crust. The latter view is summarized by Mutter and Karson (1992). Based on sea floor observations and seismic reflection images, the model involves episodes of ridge m a g m a t i s m during which complete crustal sections form alternating with episodes of tectonism during which ridge m a g m a input is low and crustal extension takes place along major lithospheric-scale faults. This tectonic rearrangement emplaces mantle material at crustal levels while largely preserving the

0040-1951/97/$17.00 © 1997 Elsevier Science B.V. All rights reserved. PII S 0 0 4 0 - 1 9 5 1 ( 9 7 ) 0 0 1 2 1 - 2

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N.H. Sleep, G.A. Barth/Tectonophysie.~ 279 (1997) 181 191

stratigraphic arrangement of mafic material above ultramafic material in each major crustal block. Shemenda and Grocholsky (1994) present a similar model based upon laboratory analog modeling in wax.

Alternatively, the model of Cannat (1993), based on sea floor and drill core observations of structure, chemistry, and stratigraphic relations among crustal components, involves the primary incorporation of mantle material into the crustal section. This model is consistent with inferences that at slow spreading ridges the lithosphere is too thick to allow long-lasting crustal magma chambers and the magma supply is too low to produce a 4 - 7 km thick mafic crustal section. The result is a model of gabbroic intrusions forming sill- or dike-like bodies surrounded by uhramafic screens which supplement the mafic material thus increasing the observed crustal thickness. The purpose of this paper is to present geophysical evidence that when formed at low spreading rates the oceanic crust contains a mixture of mafic and ultramafic components, and that this should occur even in the absence of tectonic rearrangement via extension. We provide seismic evidence consistent with the interpretation that this crust does include ultramafic material. Our thermal modeling supports the incorporation of ultramafic mantle material into the crustal section as a result of cooling from above at the ridge. We show that at spreading rates slightly too low to support a crustal magma chamber, the magma supply to the ridge is limited by freezing out of the melt at shallow depths rather than by insufficient melt production deeper in the mantle. To clarify terminology, we use the term 'crust' to mean all of the igneous and altered igneous material in the upper lithosphere that appears, via remote geophysical observations, to have lower bulk seismic velocities (<8.0 km s -~) and densities (<3.3 g cm -3) than the mantle. This may include both mantle differentiates and mantle residual material in any arrangement, proportion or state of alteration, so long as the result is geophysical crust when viewed on the scale of kilometers. 2. Seismic evidence There are several lines of seismic evidence suggesting the presence of ultramafic material in the

Atlantic V(z) data

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lower crust formed at slow (<50 mm yr -I full rate) and very slow (< 15 mm yr 1 ) spreading ridges. First, when considered on average, there is evidence for aging of lower crust formed at low spreading rates (Fig. 1). Comparing averaged seismic velocity at 1 km depth intervals for the velocity solutions compiled by White et al. (1992), there is no substantial change in lower crustal velocity with age in crust formed at spreading full-rates >70 mm yr ~ (Pacific data). But in crust formed at rates <50 mm yr I (Atlantic data) there is a systematic decrease in velocity with age, and the decrease is progressively greater with depth. The Pacific lower crustal velocities are intermediate between the young and old Atlantic values. These systematics are evident only on average; variations among individual profiles, attributable to differences in the details of crustal formation and evolution from site to site, surpass the subtle average changes with age that we report. Taking the average Pacific velocity profile as representative of a gabbroic crust, the average young

N.H. Sleep, G.A. Barth/Tectonophysics 279 (1997) 181-191

Atlantic values can be obtained by incorporation of ~ 7 - 2 0 % ultramafic material, depending upon its distribution and composition. These numbers are approximate Voigt-Reuss bounds for ultramafic compositions ranging from websterite to dunite, mixed with gabbro as sampled in Troodos hole CY-4 (Christensen and Salisbury, 1989; Carmichael, 1989). Old Atlantic values can then be reached by alteration to ~ 4 - 1 5 % serpentinite in the lower crust. Thus the change in average velocities of the low-rate crust is consistent with a plausible range of mixed mafic-ultramafic compositions. Second, comparing seismic reflection images of crust formed during slow and fast spreading, Mutter and Karson (1992) demonstrated that the Moho is a less distinct boundary and that it shows more variability in both structure and reflectivity in the slower examples. These observations are consistent with a less distinct compositional separation of mafic crust from ultramafic mantle. Also the overall reftectivity of the lower crust is much higher, including both distinct reflectors and zones of diffuse reflectivity (McCarthy et at., 1988; Hinz et al., 1989; White et al., 1990; Morris et al., 1993). Much of this reflectivity may be attributed to tectonic features such as faults and deformation zones (Mutter and Karson, 1992). It is also consistent with greater

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compositional heterogeneity relative to crust formed at higher spreading rates. The cutoff spreading rate between fast- and slow-style crustal fabric appears to be near the cutoff between axial valley and axial high ridge morphologies (Chen and Morgan, 1990). Analysis by Hinz et al. (1989) of data from the eastern Atlantic suggests an additional boundary near 20 mm yr -l full rate, with lower rates associated with formation of thicker reflection seismic crust. Third, velocity profiles of crust formed at very slow spreading ridges are distinct. There are comparatively little seismic data from crust formed at very low spreading rates. Those areas studied include the southwest Indian Ocean (Francis and Raitt, 1967; Minshull and White, 1993, 1996), the Arctic Ocean (Jackson et al., 1982, 1995), the Labrador Sea (Osier and Louden, 1992, 1995), Baffin Bay (Keen and Barrett, 1972), the Cayman Trough (Bowin, 1968), and the Porcupine, Tagus, and Iberian abyssal plains (Ginzburg et al., 1985; Pinheiro et al., 1992; Whitmarsh et al., 1993; Horsefield et al., 1994). Of these, crust of the latter five regions formed in continental rift or other narrow basin settings such that mantle source region cooling by lateral heat loss cannot be ignored (Boerner and Sclater, 1989; White et al., 1992). Crustal velocity profiles from the remaining three regions are compared in Fig. 2. While the ab-

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Fig. 2. P-wave velocity profiles through crust formed at very slow spreading rates. Arctic example is profile 3 from Jackson et al. (1982). Southwest Indian Ridge example is Lusiad 33 profile from Minshull and White (1996). Labrador Sea examples are profiles N-R2 (left) and G-RI (right) from Osler and Louden (1992). Note that all examples are younger than ~50 Ma and all were formed at < 15 mm yr -~ full spreading rate. Features making these profiles distinct from those observed at higher spreading rates include an absence of division into seismic layers 2 and 3 and an atypically high-velocity gradient through the mid- and lower crust.

184

N.H. Sleep, G.A. Barth/Tectonophysics 279 (1997) 181-191

solute velocities observed vary greatly, the profiles show two key similarities. First, there is no obvious layer 2-layer 3 boundary. Instead the velocity profiles resemble a single linear gradient from top to bottom of the crust. This is consistent with the expected absence of a crustal magma chamber which, when present, would fix the level of the top of seismic layer 3. Second, the velocity gradient through the lower crust is higher than typically observed in crust formed at higher spreading rates (compare with Fig. 1), and higher than can be accounted for by pressure-related increase in gabbroic rock velocity alone (Christensen, 1978; Christensen and Smewing, 1981; Christensen and Salisbury, 1989). The observation implies compositional change with depth in these examples. Minshull and White (1996) have found that Poisson's ratio for the southwest Indian Ocean crust is ~0.30, intermediate between values expected of gabbros and partially serpentinized peridotites. As discussed below, there is considerable similarity in seismic structure between crust formed at very low spreading rates and that observed in many Atlantic fracture zones. Based largely upon sea floor outcrop geology, Cannat et al. (1995) hypothesized that these fracture zones also contain a composite crust comprised of a mafic-ultramafic mix. Fourth, at very low spreading rates seismic crustal thicknesses appear greater than total melt produced, again suggesting that there is some supplementation of the mafic crustal material in the seismically defined crust. Global seismic velocity data compilations indicate that average oceanic crustal thickness is approximately constant for all spreading rates above ~ 2 0 - 3 0 m m yr 1 (Chen, 1992), and that it decreases significantly as spreading rate decreases below this range. Mantle melt production inferred from rare earth element (REE) concentrations in midocean ridge eruptives decreases abruptly near 1520 m m yr -] full spreading rate (Bown and White, 1994) (Fig. 3). We observe that at spreading rates >20 mm yr ', REE estimates of magmatic thickness are generally greater than or equal to seismically determined crustal thicknesses. In contrast, at lower spreading rates the seismically defined crust appears to be generally thicker than the calculated magmatic thickness. Bown and White (1994) note that due to uncertainty in both thickness determinations, individual differences are probably not significant.

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Fig. 3. Comparison of oceanic crustal thicknesses derived from seismic observations and basalt rare earth element (REE) inversion, modified from Bown and White (1994). The REE or magmatic crustal thickness appears to be systematically lower than the seismically defined crustal thickness below 420 mm yr -], consistent with our hypothesis that marie magmatic crust is supplemented by ultramafic mantle material at low spreading rates. Unfortunately, considering REE uncertainties of 1-2 km and seismic uncertainties of ~0.5 km, the absolute differences in thickness determinations cannot be used to quantify the amount of crust-mantle mixing required. 'Seismic' data points are as compiled by White et al. (1992). "Corrected seismic' points have been adjusted by Bown and White (1994) to attempt to account for variations in crustal thickness related to segment-scale magma distribution.

However we note that the change in the sign of the mismatch near ~ 2 0 m m yr -1 is in agreement with our model prediction of thickened seismically defined crust due to freezing of some marie material in the shallow mantle at very low spreading rates. Thus the available data suggest that there is incorporation of mantle material into the crustal section at both low and very low spreading rates. At low

N.H. Sleep, G.A. Barth/Tectonophysics 279 (1997) 181-191

rates, the amount of ultramafic incorporation appears to be small on average, and there is still evidence for the existence of at least an intermittent axial magma body, leading to the familiar layered crustal structure observed both seismically and in outcrop (Sinton and Detrick, 1992). In these settings, ultramafic incorporation may occur preferentially near fracture zone offsets (Cannat et al., 1995). In contrast, at very low rates evidence for a persistent crustal magma body is lacking, and some form of mafic-ultramafic mix appears to pervade the entire crustal section (Cannat, 1993).

3. Thermal modeling The thermal structure of the axial region of midoceanic ridges consists of an essentially adiabatic upwelling overlain by a shallow lid of cool material associated with the conduction of heat to the surface (and hydrothermal circulation). At high spreading rates, an upper crustal lid overlies a magma lens and a crustal mush chamber. At sufficiently low spreading rates, the cool lid extends into the mantle and a permanent crustal magma chamber does not exist. At these spreading rates, ascending melt that intrudes the cool mantle lid may freeze. This process reduces the amount of melt available to separate from the mantle and form mafic crust. It also creates a zone of intrusive mafic-ultramafic mix. As noted above, there is evidence that such a zone actually exists. It is our purpose to illustrate the physics of the formation of this crust-mantle mix. Specifically, the release of latent heat from the freezing of intruded magma increases the temperature within the cool mantle lid. The amount of melt which can freeze is limited because a maximum allowable release of latent heat would raise the temperature to the melting point and preclude further freezing of magma. We show below that, for observed spreading rates, such significant freezing of melt occurs only within a few kilometers below the depth that Moho would have at higher spreading rates. To investigate the mechanism for mantle incorporation into the crustal section, we keep the model as simple as possible. We use a steady-state two-dimensional formulation although the ridge axis is obviously three-dimensional and time-dependent when viewed in detail, We require the upwelling mantle

185

to remain as hot as possible. To do this, we make the upwelling zone very narrow, we allow enough melt to freeze at depth to maintain temperature at the melting point, and in one model we do not include heat loss from hydrothermal circulation. In a more realistic ridge model that included significant hydrothermal circulation and a broader zone of upwelling, more upwelling magma would freeze. The real mid-ocean ridge is therefore more likely to support formation of a mafic-ultramafic mix than is our simplifed model. The thermal calculations were carried out following the method of Sleep (1975) and Morton and Sleep (1985). The vertical movement of material is represented by intrusion which is assumed to occur in a very thin dike-like zone at the ridge axis so that material not exactly at the ridge axis moves at the plate velocity. The steady-state two-dimensional temperature is given by the Fourier series:

T = zTx ~ k + --

(m:rrz']

sin \ ~ - - / C m e x p ( a m )

(1)

m=l

where T is temperature, Tz -----1360°C is the boundary condition at depth X ---- 100 km (deep enough for the model to behave as a half space), and z is depth. The coefficient am is given by:

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1 --

1+

~

]

(2)

where u is the half spreading rate of the ridge, p c = 3.8 × 10 6 J m -3 °C-1 is v o l u m e s p e c i f i c heat,

and k -- 2.8 W m -l °C-l is thermal conductivity. The Fourier coefficients C,, are obtained from an energy conserving boundary condition at the ridge axis:

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dz sin(mJrz/X)S(z)

(3)

where S is the total flow of heat away from the axis:

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kOT ( T x ) o~ + u p c T - z ---~

(4)

where x is distance away from the ridge axis. In the case where pressure release melting and melt freezing occur at depth, the total heat flow is

S = upc

Ti - z--~

+ UQL + uQs

(5)

N.H. Sleep, G.A. Barth / Tectonophysics 279 (1997) 181-191

186

where Tg is the intrusion temperature at depth z, Q L is the heat extracted from latent heat by freezing of melt at depth z, and Q s is the heat extracted from superheated melt at depth z. To apply Eq. 5, it is necessary to have observational constraints or to make assumptions about the extent to which upwelling magma thermally equilibrates with the mantle. As it is our purpose to examine the effect of freezing of magmas within the mantle, we adjust QL so that the maximum amount of magma freezes at these depths. To simplify discussion of our procedure for doing this, we assume that the mantle is eutectic, having a single melting point rather than a liquidus and solidus. The physical constraint on QL is then that the release of latent heat at depth z does not raise the temperature at that point (and other points where freezing is occurring) above the melting point. At the limit which we seek, freezing of melt maintains the temperature at the melting point. We use a simplified melting-versus-depth relationship that mimics the more complex one of McKenzie and Bickle (1988). A total of 6 km of melt is produced from an adiabatic column where further melting ends at 20 km depth. The precise temperatures used are arbitrary, and can be rescaled by a constant factor. Our computational scheme keeps track of the upwelling material starting at depth Z. Below the depth at which melting starts, set at 79 km in the model, the intrusion temperature Ti decreases upward with the solid adiabatic gradient of )/ = 0.3°C km -l. Above this depth, T/ decreases upward with the melting point gradient of/3 -- 3°C km -~ . The equivalent thickness of oceanic crust M that would be formed by the melt that ascends above depth z changes as: OM - -

Oz

([3 - )/)(z - M ) p c --

L

QL +

L

(6)

where the first term is associated with adiabatic pressure-release melting, the minus sign occurs because depth is positive downwards, L = 1.026 × 109 J m -3 is latent heat per volume, and the second term is associated with latent heat released by freezing of melt at depth z. This relation was continued upward until 22% fractional melting occurred at 20 km depth. Thereafter the least abundant mineral on the eutectic was assumed to have been used up and no further

pressure-release melting occurred. The temperature T~ then decreased upward along a solid adiabat. The superheat of the magma, ~, associated with the difference between solid and liquid adiabats changes as: 0c~ Oz

-

M(/3-)/)pc+

Qs

(7)

where the second term represents heat extracted by thermal equilibration of the magma and the solid at depth z. These relationships were applied as follows using an iterative scheme. A trial temperature field was computed using Eq. 1. The intrusion temperature T/, the equivalent melt thickness M, and the superheat cr were computed. At sufficiently fast spreading rates, the computed axial temperature was greater than or equal to the melting point all the way to the model Moho, where the equivalent melt thickness was equal to the depth, M : z. The latent heat and superheat of the melt was then evenly distributed into S at crustal depths shallower than M and a final temperature field was computed. At somewhat slower spreading rates, the initial computed temperature was less than the melting point at shallow depths. A superheat source Q s was then applied to attempt to raise the computed temperature to the melting point and was adjusted using Eq. 7. If reducing cr to zero was insufficient, then a latent heat source from the freezing of melt QL was applied and the equivalent thickness M adjusted using Eq. 6. This procedure was continued upward until either a model Moho, z : M, was obtained or all the melt was used up, M : 0, where only the heat source from Ti is included in Eq. 5 above that depth. New temperature fields were computed perturbing Q s and QL until the computed axial temperature, in regions where melt was freezing, approached the melting temperature. The effect of this procedure and the approximate amount of melt that freezes at a given depth are obtained by letting the computed temperature without melt freezing be T,,. Then a local heat balance at that depth implies that: -0M*

L - Oz

~ P c ( E . -- T,:)

(8)

where - O M * / ' O z is the volume fraction of melt that freezes at depth z and Tm is the melting temperature.

N.H. Sleep, G.A. Barth/Tectonophysics 279 (1997) 181-191

Then QL = p c ( T , , - T~) would increase the temperature to the melting point if no further heat were lost by horizontal conduction in Eq. 4 from that depth. If the temperature difference in Eq. 8 is small, the perturbation of S in Eq. 5 by adding QL is small and the amount of melt which freezes is given by Eq. 8. This indicates that little melt can freeze at depths where the computed temperature without melt freezing is near the melting temperature. The amount of melt that freezes scales inversely to the latent heat L. For example, if L = 1.5 x 109 J m -3 were used (Kojitani and Akaogi, 1995), the amount of melt that freezes in the mantle would be reduced by a factor of 2/3. This procedure gives the maximum allowable freezing of melt within the mantle at the ridge axis and the highest upper mantle axial temperatures consistent with freezing of melt at depth. Other situations are conceivable. For example, all the melt formed could erupt at the surface or intrude at shallow depths without transferring any heat to the mantle. It is also conceivable that pressure-release melting is significant above the 20-kin depth where it cuts off in the model. In that case, low temperatures at the ridge axis would suppress pressure-release melting as well as freeze ascending melt. Our axial analytical boundary condition does not allow representation of hydrothermal circulation which penetrates deeply into the crust or mantle. The actual distribution of water flow in at least two dimensions needs to be included in such a model. Instead, the effect of shallow hydrothermal circulation is included in one model by setting all heat sources in the upper 2 km to zero (by including equivalent heat sinks). Computed thicknesses of melt M reaching a given depth are shown in Fig. 4. A rapid transition occurs from low spreading rates where the mantle is nearly adiabatic and a small crustal magma chamber exists to very low spreading rates where the ascending melt freezes at depth. Most of the upwelling melt freezes within ~ 1 0 km of the sea floor. A geophysically distinguishable crustal section containing some mafic-ultramafic mix is thus predicted. Temperature at the axis is shown in Fig. 5 for models at transitional spreading rates. Melt is emplaced as dikes above the model Moho and some melt freezes beneath it, mainly above 6 km depth. It is difficult to obtain a stable solution at transitional spreading

187

Melt thickness, km

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Fig. 4. Thermal models of slow spreading ridges, after Sleep (1975). Adiabatic crustal thickness (at high spreading rates) is 6 km for the case modeled. Contour lines are available melt in the column M (i.e., melt produced and not subsequently frozen out), in km. Model Moho (M) is drawn where the distance to the surface is equal to the amount of melt remaining in the column, so that it marks the maximum possible marie crustal thickness. Below the critical spreading rate, some of the melt produced at depth is frozen into the uppermost mantle rather than successfully separating from the mantle material. The geophysically detected Moho is likely to be at the base of this region of crustmantle mix. Note that although it cannot successfully arrive at the surface, adequate melt is produced at depth to form a full 6 km thick marie section at these spreading rates.

rates, because the intrinsic problem is unstable. That is, a small decrease in spreading rate reduces the melt available to form model crust and thus reduces the temperature at the base of that crust to where latent heat cannot maintain temperature at the melting point. A possible solution was approached by iteration. This tended to smooth the axial geotherm and make the model Moho temperature differ from the melting point, as in Fig. 5. As with other thermal models, a low-temperature lid occurs above an essentially adiabatic region at the ridge axis. Within the adiabatic region geotherms are the same as those

N.H. Sleep, G.A. Barth/Tectonophysics 279 (1997) 181 191

188

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Fig. 5. Thermal models of the axial region at spreading lull-rates 13 mm yr -I neglecting hydrothermal circulation and 14 mm yr-] including hydrothermal circulation. These are spreading rates where surface cooling has limited the amount of melt available at the surface despite adequate melt production at depth. Note that freezing occurs within 10 km of the sea floor. In each case enough heat is liberated by melt freezing to keep the mantle at the eutectic temperature beneath the ridge. The model Moho does not exactly converge to the top of the adiabatic zone because the solution method is damped somewhat for numerical stability.

from a model where broad upwelling occurs (mathematically the gradient of temperature is zero there so the exact arrangement of velocity does not enter into the heat flow equation) (e.g., Sleep, 1969). The critical spreading rate (below which melt is incompletely separated from the mantle) in the absence of hydrothermal cooling scales inversely with the thickness of melt produced in an adiabatic column. Neglecting hydrothermal cooling, for 6 km adiabatic crustal thickness, 13.4 m m yr -] (full-rate) is the calculated critical spreading rate. The critical rate is higher, 14.2 m m yr -I, when hydrothermal cooling is included. Cooling depth is inversely proportional to spreading rate, so for a given melt supply a slower ridge will produce a thicker section of maficultramafic mix. In the case modeled, deep cooling begins to severely reduce the amount of melt produced near 8 mm yr -I spreading full-rate. We stress that we have generated a fairly simple model which can sustain normal mafic crustal production to lower spreading rates than may be realistic. A finite width of mantle upwelling would affect the deep cooling, shifting the melt production drop-off toward higher spreading rates. Adding more extensive hydrothermal circulation and broadening the zone of crustal injection would enhance shallow cooling, shifting the critical spreading rate toward higher values as well. A well-known effect of low spreading rate and consequent deep cooling is a reduced amount of

melting in the mantle source region (e.g., Bown and White, 1994). Our two-dimensional steady-state models indicate that only the region above about 10 km depth is cooled at observed spreading rates. Our source region which is below 20 km depth is not affected. In general, cooling from the surface causes freezing of magma to become significant at higher spreading rates than those needed to significantly reduce melt production at depth. Two-dimensional cooling of crust and uppermost mantle at low spreading rates, as modeled here, is distinguished from three-dimensional deep cooling of the source region at long transform offsets, after continent breakup, or after ridge jumps. These special cases will lead to reduced melt production at higher spreading rates than the two-dimensional heat flow which we have modeled will. The model results imply that at spreading rates too low to sustain a steady-state magma chamber, some degree of mafic-ultramafic mixing should be occurring due to cooling at the ridge. In the absence of other factors that may alter crustal thickness (discussed below), the model predicts a systematic increase of the geophysically determined crustal thickness as spreading rate decreases below the critical rate, due to incorporation of mantle material into the crustal section. This trend should be reversed at very low rates where source region cooling diminishes melt production. At spreading rates below the crit-

N.H. Sleep, G.A. Barth/Tectonophysics 279 (1997) 181-191

ical rate for complete separation, the geophysically observed crustal thickness should exceed the total thickness of melt produced (with the caveat that melt freezing at substantial depth may not be seen as part of the geophysical crust).

4. Discussion Our thermal model is consistent with seismic evidence for incorporation of some mantle material into the oceanic crust, and for increasing amounts of mantle involvement as spreading rate decreases. However, the model does not predict the inferred onset of mantle incorporation at spreading rates as high as 50 m m yr - t . Also it does not simply predict the abrupt drop-off of crustal thickness and melt thickness near 15-20 m m yr -1. These differences are likely due to two simplifying assumptions incorporated into our model. One is the assumption of continual dike intrusion to sustain some form of axial magma body to the lowest possible spreading rates. The other is the assumption that narrow dike-like axial flow will continue in the mantle at all spreading rates. We have modeled the slow-spreading ridge axis including continual dike intrusion into a very narrow axial zone in the lower crust. The minimum crustal magma chamber in the models is one where the very deepest part of a dike is just about to freeze when the next dike comes in. In reality, a larger axial magma chamber is necessary to be detectable and have significant mechanical effects. The axial magma input is likely to be episodic rather than continual and the freezing of melt to maintain mantle temperatures at the melting point is not likely to be perfectly efficient, implying that the ideal conduit allowing maximum melt-mantle separation is unlikely. Instead we envision a less well organized process of melt arrival and freezing beneath the spreading center, leading to the onset of incomplete separation at higher spreading rates. Thus at spreading rates somewhat above our critical full-rate (i.e., >15 m m yr -1) primary mafic-ultramafic mixing must already be occurring. Our models assume that intrusion occurs in a narrow region at the ridge axis. At great depths where the models predict a broad adiabatic zone this assumption as noted above has the same implications as a broad adiabatic upwelling. In actuality, both

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the zone of dike intrusion and the zone of mantle upwelling have finite widths. A wider intrusion zone somewhat lowers axial temperatures. (The effects of a finite zone of intrusion are approximately obtained by averaging temperatures over the zone width in Fig. 5.) The effects of a finite zone of mantle upwelling are stronger because of the strong feedback between temperature and viscosity. That is, at slow spreading rates the predicted width of the conduit for ascending mantle material is narrow and high stresses are needed for material to flow. Such stresses also cause significant deformation in the subsolidus crust and mantle effectively broadening the zone of upwelling. The broadening of upwelling further lowers axial temperatures thus favoring still broader upwelling. The limiting case of broad upwelling behaves like comer flow where the width of upwelling is comparable to depth. Bown and White (1994) have successfully reproduced the general spreadingrate dependence of mantle melting using a mantle corner-flow model including decompression melting but neglecting latent heat and hydrothermal cooling. There is considerable similarity in seismic structure between crust formed at very low spreading rates and that observed in Atlantic fracture zones (Pinheiro et al., 1992; Minshull and White, 1996). In both settings the abnormal upper crustal gradients and lack of prominent layer boundaries have been attributed to possible pervasive fracturing of crust, during amagmatic extension or transform slip (Osier and Louden, 1992; Detrick et al., 1993). This tectonization and related alteration should result in lowered velocities in the mafic crust. However, in several low-spreading rate examples lower crustal velocities are quite high (Fig. 2). Also, in laboratory experiments crack closing has already been completed by 2 kbar pressure ('-qithostatic at base of crust), making pervasive fractures alone a poor explanation for velocity gradients in the lower crust. Cannat et al. (1995) have proposed that many Atlantic fracture zones contain a mixed mafic-ultramafic composition, akin to that proposed for very low spreading rates. Tectonization and associated serpentinization of this mixed composition has more potential to cause the large velocity variations observed. In addition, much of the material identified as partially serpentinized peridotite in fracture zones and beneath thin crust may actually be a mafic-ultramafic

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mixture, thus reducing the need for pervasive mantle alteration to explain the geophysical observations. Typically a seismic Moho will be interpreted to exist at the base of a region of mixed maficultramafic composition. As we have argued, the extent of mantle incorporation into the crustal section may have an important effect on observed crustal thickness at very low spreading rates (Mutter and Karson, 1992; Cannat, 1993; Cannat et al., 1995). However, a variety of factors simultaneously affect the thickness of oceanic crust produced at mid-ocean ridges. These include mantle temperature differences, which lead to differences in extent of melting (Klein and Langmuir, 1987; Bown and White, 1994), upwelling pattern and extent of alongaxis flow, which affect magma distribution (Lin and Phipps Morgan, 1992), and tectonic extension, which may alter the original magmatic thickness of crust during or after emplacement (Mutter and Karson, 1992; Escartin and Lin, 1995). As a result, the contribution of primary mafic-ultramafic mixing may be difficult to distinguish. Crustal thickness measurements aside, we expect that the presence of the predicted mixed composition can be identified on the basis of velocity structure and Poisson's ratio. Young crust of mixed composition is expected to exhibit unusually high velocities and gradients in the lower crust. Where possible, serpentinization may systematically reduce lower crustal velocities and velocity contrasts while raising Poisson's ratios, thus the seismic signature of a mixed composition may vary markedly as the crust ages.

5. Conclusions When produced at low (<50 mm yr - I ) and very low (<15 m m yr - l ) spreading rates, the geophysically identified oceanic crust appears to contain significant amounts of ultramafic mantle material. At low spreading rates this is evidenced by sea floor observations, aging of the lower crust seen as reduction in average seismic velocity, and a more heterogeneous lower crust and more diffuse Moho boundary as seen in seismic reflection images. Our modeling indicates that the incorporation of mantle material may be attributed to tectonic rearrangement at the higher spreading rates (Mutter and Karson, 1992),

but that at rates approaching the very low end of the spectrum there is almost certainly some primary mixing of crust and mantle material due to incomplete separation of melt as it ascends beneath the cool ridge. At very low spreading rates the presence of a mixed composition is evidenced by sea floor observations, atypically rapid increase in seismic velocity with depth in the lower crust, the absence of a layer 2-layer 3 boundary, and the suggestion that geophysically observed crustal thickness is greater than the total thickness of melt produced. Our modeling indicates that at these very low spreading rates, incomplete separation of mafic melt from ultramafic mantle occurs as a consequence of cooling from above, and that there is no need to call upon tectonic rearrangement of a normally stratified oceanic lithosphere to explain the observations.

Acknowledgements We acknowledge M. Cannat and two anonymous reviewers, and thank M. Cannat, J. Osier and T. Minshull for providing preprints as well as enlightening discussions. This work was supported in part by NSF grants OCE-9402799, OCE-9302184, and EAR-9204708.

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