Geophys. J. R. astr. Soe. (1 978) 54,63 1-660
The Mid-Atlantic Ridge at 37 and 45" N: some
geophysical and petrological constraints
EUm G . Nisbet* Institut fur Kristallogrophie und Petrogrophie, und
Geologisches Institut, ETH, CH 8092 Zurich, Switzerland
c. Mary R. Fowler t Institut fur Geophysik, ETH-Honggerberg,
CH 8093 Zurich, Switzerland
Received 1978 February 1; in original form 1977 September 19
Summary. At present there is a strong conflict between, on the one hand,
seismological and thermal models of the Mid-Atlantic Ridge, which indicate
that no large crustal magma chamber can exist, and on the other hand petrological models many of which stress the importance of such a chamber. We
review the available geophysical and petrological information from the
FAMOUS area and 45" N in an attempt to resolve this conflict and demonstrate that a model (the infinite leek) can be constructed which satisfies all
the available seismological, thermal, petrographic, major element and traceelement information from these two areas. This mode is as follows: mantle
rising from depth begins to melt at about 60 km, and rises in equilibrium with
its melt to about 15-25 km below the sea surface. At this level melt
segregates and rises rapidly to the base of the crust. Magma injection above
this takes place by a process of crack propagation, or by the development
of a narrow vertical magma chamber, but no large crustal chamber is present.
This model successfully explains the marked petrographic zonation of the
floor of the median valley (Hekinian, Moore ck. Bryan).
1 Introduction
In recent years many scientists have studied the Mid-Atlantic Ridge, particularly in the
FAMOUS area (37"N) and also near 45"N. Many petrological workers7e.g. Cann 1974;
Bryan & Moore 1977; Ballard & van Andel 1977) have emphasized the importance of a
supposed large crustal magma chamber under the whole inner floor of the median valley,
while seismic data ( e g . Fowler 1976, 1978; Steinmetz, Whitmarsh & Moreira 1977) have
indicated that such a magma chamber does not at present exist, though of course it may
have existed in the past. The purpose of this paper is t o re-examine both the geophysical
and petrological constraints on the structure of the Mid-Atlantic Ridge, in the hope that a
model can be constructed that is consistent with all the available information, and which
resolves the present conflict between seismic and petrological evidence.
*Present address: Department of Mineralogy and Petrology, Downing Place, Cambridge.
t Present address: Department of Geodesy and Geophysics, Madingley Rise, Cambridge.
632
E. G . Nisbet and C. M.R . Fowler
The discussion in this paper is confined to the Mid-Atlantic Ridge in the two intensively
studied areas 37 and 45" N. We shall not make extensive comparisons between the MAR and
the East Pacific Rise, another intensively studied segment of ridge; the two ridges have
different surface topography and probably also different seismic structures. Comparison of
seismic record sections from axial refraction lines shot at 37" N on the MAR (Fowler 1976,
Fig. 12) and 9"N on the EPR (Orcutt, Kennett & Dorman 1975, Fig. 2) suggests that the
velocity/depth structures of the two areas must be quite different.
In the following discussion we shall separately set out the geophysical and petrological
information on the structure of the MAR. We shall then attempt to build up models which
are consistent with all the available evidence. The analogy with ophiolite complexes provides
a useful, though not definitive, test for any model of the oceanic crust. Since the paper is
interdisciplinary and complex, we have attempted at each stage of the discussion to provide
summaries of the conclusions reached.
2 Geophysical constraints on the structure of the MAR
2.1 S E I S M I C R E S U L T S
In the last few years there have been several seismic experiments on the MAR at 37 and
45"N. These are described by Keen & Tramontini (1970); Whitmarsh (1973); Fowler &
Matthews (1974); Poehls (1974); Whitmarsh (1975); Fowler (1976); Steinmetz et al. (1976,
1977) and Fowler (1978). In addition to these seismic experiments, microearthquake
studies have been carried out by Francis & Porter (1973); Reid & MacDonald (1973);
Spindel et af. (1974); Francis, Porter & McGrath (1977) and Lilwall, Francis & Porter
(1977).
Whitmarsh (1973), Fowler & Matthews (1974) and Poehls (1974) used first arrival times
to model the seismic structure of the ridge crest region in the FAMOUS area; these three
experiments yielded three rather different models. In Fowler (1 976) these three interpretations were re-examined and a more detailed model for the seismic structure of the ridge was
proposed ; this model was derived by comparing synthetic seismograms computed from trial
structures with the observed seismograms, and so eliminating incorrect models. The final
model proposed by Fowler (1976) for the FAMOUS area produces synthetic seismograms
which match the observed seismograms well (Fowler 1976, Figs 12, 14 and 16). This model
is shown in Fig. 1. The model has material with a normal upper-mantle velocity present close
to the ridge axis, while actually at the ridge axis the upper-mantle velocity is low. The crustal
structure on the ridge flanks is very similar to the normal oceanic crustal structure (but with
layer 3 split into 3A and 3B), while the crust at the axis is thinner and there is no normal
crust-upper mantle discontinuity at the axis. In particular no evidence was found for any
crustal low-velocity zone along the ridge axis. Work by Keen & Tramontini (1970), Fowler
(1978) and Fowler & Keen (in preparation) has demonstrated that, within the limits of the
available data, the structure at 45" N is very similar to that at 37" N. Steinmetz et al. (1977)
have investigated the amplitudes of first arrivals from a refraction line crossing the MAR
axis at 40" N; they also concluded that their data did not indicate the presence of a crustal
magma chamber. Whitmarsh (1975) suggested that there is a small low-velocity zone at the
sea-bed along the axis of the MAR in the FAMOUS area (this is discussed further in Fowler
1976). This would seem to be a shallow zone associated with the extrusion of lavas at the
sea-bed. In contrast, the crustal low-velocity zones which have been proposed to exist along
the axis of the East Pacific Rise (Orcutt et al. 1975 ;Rosendahl et al. 1976) are generally
assumed to be magma chambers similar to the 'infinite onion' proposed on geological
grounds by Cann (1968, 1970, 1974). The presence of these low-velocity zones has been
deduced from P wave travel times and from amplitudes.
Mid-Atlantic Ridge at 37and 45" N
633
kin
30
20
10
0
10
0
52
E
Y
.
66
72
..=fig'-...._ /:..
-*
I
,-
"
-
20
30
sea
layer 2
layer 3A
layer 38
mantle
7.6
Figure 1. Seismicvelocity model for the MAR crest in the FAMOUS area, after Fowler (1976). Stippled
region - extent o f magma chamber proposed on petrological evidence (Bryan & Moore 1977; Hekinian
et al. 1976). Solid region - extent of steady-state magma chamber possible on thermal evidence (Sleep
1975). Numbers: P-wave velocity in km/s.
Perhaps the most conclusive test of the presence of any partially molten zone can be
provided by shear waves: shear waves crossing such a zone should either be greatly
attenuated or absorbed. An array of three-component seismographs has been used to
monitor the microearthquake activity at 37 and 45" N on the MAR by Francis et al. (1 977)
and Lilwall e t al. (1977). These same instruments were used to record the shots fired in the
refraction experiments reported by Fowler & Matthews (1974) and Fowler (1976, 1978).
With a three-component instrument positive identification of shear-wave phases is possible.
The microearthquakes recorded by the OBS both at 37 and 45" N all show large-amplitude
shear waves and there is no evidence for any special crustal attenuation of these phases.
For the refraction lines shot at 37 and 45" N shear waves were recorded from almost every
shot over 100 lb with a range greater than 10 km and can be positively identified. Amplitude
measurements of these shear waves show no indication of any crustal absorptive zone. The
crustal shear waves have travel times consistent with their having crossed the axial zone at
depths down to about 5 km beneath the sea-bed. At these depths they would have had
a wavelength of up to 1 km. Reid & MacDonald (1973) and Spindel et al. (1974) reported
shear waves from median valley and fracture-zone earthquakes which occurred during
their hydrophone-array microearthquake surveys in the FAMOUS area. All this shear-wave
evidence strongly suggests that no large crustal magma chamber can be present beneath
the median valley of the MAR either at 37 or at 45" N. Small pockets of melt of the order
of up to perhaps 2 km in diameter cannot be excluded, but a magma chamber the width of
the median valley as proposed by Bryan & Moore (1977), Hekinian, Moore & Bryan (1976)
and others, cannot easily be reconciled with the seismic evidence.
The structure of the upper mantle in this region has been studied recently by Steinmetz
et al. (1976, 1977) and Fowler (1978) by shooting long refraction lines parallel and
perpendicular to the ridge axis. These experiments suggest that there must be an absorptive
(low-velocity and/or low-Q) zone in the upper mantle, along the ridge axis, deeper than
about 6-7 km beneath the sea-bed. The width of this zone is unknown, but it must be in
excess of 20 km or so at depth. Solomon (1973) and Solomon & Julian (1974) also deduced
the presence of a low-velocity/low-Q zone in the upper mantle, beneath the MAR, probably
confined to depths between about 10 and 50 km beneath the sea-bed at the ridge axis and
with half-width several tens of kilometres. Forsyth (1975, 1977) has investigated the shearwave structure of the mantle in the Pacific using surface-wave dispersion analysis on teleseismic events. This work has indicated the presence of a zone up to about 80 km in
half-width between the depths of roughly 30 and 6 0 km beneath the sea-surface at the ridge
634
E. G.Nisbet and C, M.R . Fowler
axis in which the mantle shear velocity is reduced. Forsyth interprets this zone as a region
of extensive melting. The Pacific surface-wave work of Leeds, Knopoff & Kausel (1974)
and Yoshii (1975) also suggests that near the ridge axis the lithosphere is only 10 or 20 km
thick. Weidner (1 974) investigated surface-wave phase velocities in the Atlantic and found
that 'phase velocities deduced for the ridge imply a shear velocity which is significantly
lower than that for the ocean basins for depths in excess of 20 km to at least 200 km'.
Forsyth (1975) notes that deeper than about 80 km there is no significant change in the
structure of the mantle over the first 40 Myr. The bottom of the low-velocity zone in this
region is found to be somewhere between 150 and 200 km (Forsyth 1975; Leeds 1975;
Yoshii 1975; Weidner 1974). This depth is typical for the base of the oceanic asthenosphere
(low-velocity zone), but the average shear velocity of the asthenosphere apparently increases
from about 4.1 km/s under young sea-floor (- 40 Myr) to 4.25 km/s under old sea-floor
(- 120 Myr) (Yoshii 1975; Forsyth 1977).
2.2
OTHER GEOPHYSICAL RESULTS
There are two recent thermal models of the formation of the oceanic crust, the first by Sleep
(1975) and the second by Kusznir & Bott (1976). The conclusions reached by these two
studies are broadly the same; Sleep concluded that for mid-ocean ridges with a halfspreading rate less than 0.9 cm/yr the latent heat of basalt in the top 5 km of the oceanic
crust is insufficient to maintain a steady-state magma chamber, while Kusznir & Bott concluded that no steady-state magma chamber could form for half-spreading rates less than
0.45 cm/yr.
The approaches used in these two studies are basically the same, the main difference
being in the boundary conditions placed on the differential equation: Kusznir & Bott
following McKenzie (1967) specified a constant temperature along the intrusion axis, while
Sleep, considering adiabatic and melting gradients, specified the heat flux along this axis.
Kusznir & Bott's assumption of a constant temperature along the intrusion axis results in a
singularity in the heat flow at the axis, while Sleep assumed that all the latent heat is
effectively released at the axis. Sleep's assumptions meant that he was able to find
an analytical solution, while Kusznir & Bott (solving the problem by a finite difference
approximation) were able to investigate the difference between cases in which the latent
heat is released between the liquidus-solidus temperatures and those in which it is released
entirely at the solidus temperature.
For a half-spreading rate of 1 cm/yr Sleep's model yields a magma chamber with a
maximum half-width of about 0.5 km, while Kusznir & Bott's model yields a chamber with
a maximum half-width of about 3 or 4 km (fo;latent heat released entirely at the solidus or
distributed over the liquidus-solidus i n t d a l , respectively).
Some of this apparent difference between the two models is due to the different temperatures used in the two studies and some due to the different treatments of latent heat. Sleep
defined the edge of the magma chamber by the 1185°C isotherm and assumed a temperature
of 1290°C at the base of the lithosphere, while Kusznir & Bott defined the edge of the
magma chamber by the 1000°C isotherm and assumed an intrusion and base of the lithosphere temperature of 1 125°C. An estimate of 1203-1246°C for the eruption temperature
of basalts collected from the FAMOUS area has been made by Hekinian et al. (1976). Tilley ,
Yoder & Schairer (1964) reported the melting of a MAR basaltic glass: they noted the disappearance of clinopyroxene at 1165"C, plagioclase at 1215°C and olivine at 1245°C.
Kushiro (1973) has reported the melting of a MAR olivine tholeiite, yielding the disappearance of clinopyroxene at 1 160"C, plagioclase at 1180°C and olivine at 1210°C (extrapolated
Mid-Atlantic Ridge at 37and 45" N
635
for pressures of 1.5 kb). Since these are the important phases involved in crystallization in
the magma chamber, temperatures in this range would seem to be appropriate to the thermal
analysis. It would thus appear that Sleep's temperature assumptions are more realistic than
those made by Kusznir & Bott. We have therefore recalculated the thermal model of Kusznir
& Bott using liquidus temperatures of 1230-1250°C and solidification temperatures of
1160-1 185°C. For these models and a half-spreading rate of 1 cm/yr the half-width of the
magma chamber is about 2 km, (this would be reduced if the latent heat were released
entirely at the solidus).
Kusznir & Bott considered the important effect of accumulation of crystals settling to
the bottom of the magma chamber (see later). They found that a 2 : l ratio of solidified
material forming the chamber roof to that falling to the bottom reduces the maximum
width of the magma chamber to about one-half of its previous value, while a 1 : 1 ratio
reduces the maximum width to about one-third of its previous value. It seems, therefore,
that using the experimental melting temperatures and considering crystal settling for halfspreading rates of about 1 cm/yr, these thermal models of Sleep and Kusznir & Bott suggest
that the maximum half-width of any steady-state magma chamber is of the order 0.51O
. km only.
One point considered by neither of these studies is the presence of possible xenocrysts
in many of the basalts erupted in the FAMOUS area and at 45" N (this is discussed in Section
4.3). These crystals typically show the effects of resorbtion (reaction with the melt); thus
they may represent an important heat sink in the postulated magma chamber as they melt.
Such a heat sink would have the effect of increasing the minimum spreading rate at which a
steady-state magma chamber can form. A second important point is that recent work in the
FAMOUS area (Arcyana 1975; Bryan & Moore 1977) has suggested that flank eruption at
the edge of the median valley may be a small but important process in the generation of new
oceanic crust. Thus it is possible that some contribution to the net spreading rate of the
crust may come from the edge rather than the centre of the valley; the material at the centre
may be spreading at a lower rate (Deffeyes 1970). This would have the effect of increasing
the minimum spreading rate at which a steady-state magma chamber can form.
To summarize : thermal considerations imply that a steady-state crustal magma chamber
can exist in fast spreading ridges, but for half-spreading rates less than 1 cm/yr it is questionable whether any steady-state crustal magma chamber can exist. For the FAMOUS area
(half-spreading rate - 1.1 cm/yr) the maximum half-width of such a chamber is probably
0.5-1 .O km. At this spreading rate any chamber would be confined to depths equivalent to
the lower part of layer 3 (Fig. 1). There have been suggestions that spreading in the
FAMOUS area is asymmetric (0.7 and 1.5 cm/yr (Needham & Francheteau 1974) and 1.0
and 1.2 cm/yr (Greenwalt & Taylor 1974)). Sleep (1975, Fig. 3) has calculated the size of
the steady-state magma chamber for the rates of Needham & Francheteau: the resulting
chamber is slightly asymmetric but essentially not different in width from the chamber for
a spreading rate of 1.1 cm/yr.
3 Petrological constraints
It has been known for many years that most ocean-ridge basalts are remarkably similar
chemically. The vast majority contain olivine, hypersthene, diopside and plagioclase in the
norm (that is they are olivine tholeiites); they may contain phenocrysts of olivine or plagioclase, rarely clinopyroxene but never orthopyroxene; typically have a fine-grained groundmass, often glassy or cryptocrystalline; and appear to have been extruded as sea-bottom lava
flows (although intrusive bodies are also widely reported as a component of the oceanic
E. G. Nisbet and C. M.R. Fowler
636
upper crust). (The norm is the theoretical mineral assemblage of a rock expressed in terms of
standard minerals, and calculated according to a set procedure. It is widely used by igneous
petrologists for classification, and in basalts gives a useful notion of the degree of silica
saturation. Tholeiitic basalts contain normative quartz (Q) and hypersthene (hy), olivine
tholeiites (typical ocean-ridge basalts) contain normative olivine (01) and hypersthene, and
alkali basalts contain normative olivine and nepheline (ne), in addition to clinopyroxene
(diopside) and plagioclase (anorthite) which are common to all three types.) This similarity
between the vast majority of ocean-ridge basalts provides some of the most important
petrological constraints on the genesis of oceanic crust. Recent intensive work in the
FAMOUS area has shown that much more geochemical variation exists on the ocean floor
than had previously been suspected, but nevertheless has confirmed that virtually all oceanfloor lavas lie within a closely-bound compositional field.
The composition of basalts is generally expressed in terms of a series of major elements,
quoted as oxides, and trace elements. O’Hara (1968a’b) and Jamieson (1969, 1970)
suggested examining the petrogenesis of oceanic lavas by using ‘CMAS’ projections. In these
projections the composition of the rock is reduced to four components: ‘C’ = XO = molecular proportions (CaO - 31/3P20, t 2Na20 t 2 K 2 0 ) x 56.08; ‘My = YO = molecular proportions (FeO t MnO + NiO t MgO - TiOz) x 40.31 ; ‘A’= R203=molecular proportions
(AlZO3tCr203+ Fe203+Naz0+ K 2 0 + TiOz) x 101.96; ‘S’= Z 0 2 = molecular proportions
(Si02 - 2Na20 - 2K20) x 60.09. Any basalt can be plotted in the CMAS tetrahedron, as can
key minerals such as olivine (M2S), orthopyroxene (MS), basaltic clinopyroxene (CMS2) and
plagioclase (CAS2). Phase boundaries determined by experiment can also be plotted in this
tetrahedron. Phase boundaries used here (Fig. 2) are from O’Hara (1968b), derived from
,experiments by several authors (mainly on Hawaiian and mid-ocean basalts). The boundaries
are not necessarily correct for the Mid-Atlantic Ridge basalts, but they probably represent
an approximation to the truth. The phase boundaries define a number of phase volumes in
the CMAS tetrahedron - for example, if a basaltic liquid is in the olivine phase volume it
will precipitate olivine and itself move away from olivine in composition until it reaches the
next phase boundary away from olivine, when precipitation of another phase may begin.
Since the CMAS tetrahedron is three-dimensional (itself a projection of 13 dimensions) it
is more convenient to study sections through the tetrahedron. Three projections are
commonly used: from diopside (clinopyroxene), from olivine and from enstatite (orthopyroxene) (Fig. 2).
It should be noted that these methods have certain severe drawbacks - the projection
gives no information about the relative proportions of Fe and Mg, and the phase boundaries
are sensitive to variation in important minor components, such as Ti02 and HzO. Nevertheless they constitute one of the most powerful means of depicting the controls on basalt
chemistry.
3.1
CONTROLS ON BASALT CHEMISTRY
Two distinct lines of thought have governed the interpretation of the observed chemical
variation in oceanic basalts: analogy with experimental work on partial melting of probable
mantle source rocks, and modelling of fractional crystallization trends which are postulated
to control element distribution.
Some major controls on the composition of ocean-floor basalts were outlined by OWira
(1968a). Most ocean-ridge basalts appear to lie between two important barriers or ‘thermal
divides’. These divides are illustrated in Fig. 2(a). The ‘olivine-gabbro’ divide inhibits
evolution between quartz and olivine normative liquids until pressures less than about 5 kb
63 7
Mid-Atlantic Ridge at 37 and 45” N
A
5).
9
,
Orthopyroxene
Ollvlne
LO
C 3 A = LO
30
20
10
20
30
LO
50
M
i
D i o p s i d e 50
tcs-
LO
30
20
10
MS
Enstotile
\
t
l
Figure 2. Outline controls on melting and crystallization, after O’Hara (1968a). Field includes majority
of ocean-floor basalts analysed from MAR. Isobaric phase boundaries and divariant phase asseniblages
constructed from basalt melting relations. Pressures on isobaric boundaries in kilobars, except 1 atmosphere boundary. Isobaric boundaries in the diagrams show the nature of the phase expected to
precipitate from a liquid together with the ‘projection’ phase. Liquids evolve away from the precipitating
phases. In the melting of peridotite, initial liquids are produced a t points similar to ‘L’ (for each pressure),
in equilibrium with phases shown and projection phase. As melting proceeds, liquids move along
boundary between remaining residual phases. For detailed description see O’Hara (1968a, b). (a) Projection from diopside CMS, into the‘ plane M-S-C,A. (b) Projection from olivine M,S into the plane
P1 = plagoclase, 01 =
CS-MS-A. (c) Projection from enstatite MS into the plane M,S-C,S,-A,S,
olivine, GI = garnet, Cpx = clinopyroxene, Opx = orthopyroxene, Sp = spinel. Pyrope is a garnet, diopside
a clinopyroxene, enstatite an orthopyroxene.
638
E. G. Nisbet and C. M.R.Fowler
(= depths shallower than about 15 km, 1 kb being very roughly the pressure exerted by 3 km
of ocean-floor rock), and the ‘hypersthene-gabbro’ divide inhibits evolution between
hypersthene and nepheline normative liquids at intermediate pressures.
In addition, O’Hara (1 968a) suggested that to produce ocean-ridge basalts from a
peridotitic source rock (a favourite candidate for parent mantle, containing olivine, orthopyroxene, clinopyroxene and plagioclase, spinel or garnet) by equilibrium partial melting,
the separation of melt from parent must have taken place either at (a) depths between 15
to 20 or 25 km, or (b) deeper than 75-90 km. An initial liquid produced in either of these
depth ranges would be ol-hy normative (see Fig. 2), and could rise to the surface, precipitating olivine as it rose and itself moving away from olivine in composition. It would
arrive at the surface as an olivine tholeiite. O’Hara’s observations were made when only a
very limited number of specimens of ocean-ridge basalt were available, but the general
constraints of his scheme still hold although some quartz and nepheline normative oceanfloor rocks are now known (e.g. Kay, Hubbard & Gast 1970, and see later in this paper).
Of the two alternatives O’Hara favoured (b), deriving liquids from deeper than 75-90 km.
It should be noted that the geophysical work discussed above (e.g. Solomon 1973; Kausel
1972 -quoted in Forsyth 1977; Solomon & Julian 1974) makes this unlikely as the
anomalous low-velocity zone beneath the ridge extends down to perhaps 60 km, rather than
beginning at 75-90 km and extending below it, as implied by hypothesis (b).
In contrast to O’Hara, Green & Ringwood (1967), Green (1971) and Green & Lieberman
(1976) have constructed a model of ocean-ridge vulcanism based on ‘pyrolite’ (a hypothetical chemical mixture of one part basalt and three parts peridotite). Green (1971, Fig. 3)
showed that plagioclase (an aluminous mineral) would be preferentially eliminated in the
early stages of partial melting of such a mixture at pressures up to about 10 kb. Thus
aluminous olivine tholeiite melts can be produced at these depths. Between 10-20 kb
preferential melting of aluminous clinopyroxene is important. Initial melts would be
nepheline normative (see Fig. 2a) but as melting progresses and clinopyroxene is eliminated, orthopyroxene in increasingly drawn into the melt and the liquid becomes hypersthene
normative, giving a less aluminous parental olivine tholeiite. Green & Lieberman (1976)
concluded that %igh-alumina’ (greater than 15 per cent Al,O,) oceanic basalt liquids are the
result of 20-25 per cent partial melting of source pyrolite at about 10-15 kb, while ‘lowalumina’ (less than 14 per cent A1203)
parental liquids are derived by 25-30 per cent partial
melting at 15-20 kb. These liquids may then fractionate (mainly by precipitating olivine)
on ascent to give the observed ocean-ridge basalts.
There are thus three alternative sources proposed for the parental liquids of ocean ridge
basalts :
(1) Separation of the melt from residual peridotite at great depth (75-90km or deeper),
followed by rise of the melt, precipitating olivine on ascent. The degree of melting would
probably be fairly small. This model is unlikely, if only because of the geophysical evidence
discussed above.
( 2 ) Separation of the melt from the residue in the range 30-60 km, with higher degrees of
melt (c. 20-30 per cent). This process would produce low-alumina parent liquids if separation took place at around 45-60 km, leaving an olivine-orthopyroxene residue; it would
produce high-alumina parent liquids if separation took place around 3 0 - 4 5 km,leaving an
olivine-orthopyroxene-minorclinopyroxene residue.
(3) Shallower melting, with final separation of the melt from parent at less than 25 km,
but deeper than c. 15 km, to give olivine tholeiites. This latter model is supported by
Kushiro & Thompson (1 972) and Kushiro (1 973).
Mid-Atlantic Ridge at 37and 45"N
639
It should be pointed out that the range of proposed solutions to this problem is so great
that it has driven at least one noted authority to Popper (Carmichael, Turner & Verhoogen
1974, p. 628). We have not discussed the significance of water in the problem (see Kay et
al. 1970); however, Bryan & Moore (1977) noted that the water content of primitive
erupted basaltic liquids is very low (in one sample (528) HzO+ is 0.12 per cent in a ground
powder) suggesting that the source is fairly dry. Kushiro (1973) noted that some liquids
may be contaminated on ascent by water from altered wall rocks; this may affect high level
fractionation processes.
Much other recent petrological work has concentrated on the problems of fractional
crystallization at high level, in an attempt to identify parental liquids and to relate analysed
samples to one another by fractionation trends (e.g. Blanchard et al. 1976; Bryan & Moore
1977). Trace-element work (e.g. Frey, Bryan & Thompson 1974) has also attempted to
identify specific minerals involved in fractionation, as well as in partial melting.
At this stage the more geophysically inclined reader may wish to inspect Figs 3 and 4 and
then pass on to the 'Summary of petrological constraints'.
4 Evidence from basalts from DSDP Leg 37, the FAMOUS area and 45" N
A very large number of basalts has been collected from the FAMOUS area by dredging
and diving (Hekinian et al. 1976; Bryan & Moore 1977). DSDP Leg 37 cored many samples
from 37" N very close to the FAMOUS area and on a stream line from it. In the following
discussion we re-examine published evidence from these areas in an attempt to determine
the conditions of melting and fractional crystallization. In some cases our interpretations
conflict radically with those of the authors we have cited; we have based our conclusions
entirely on available published data, both geochemical and petrographic, not on published
interpretations.
4.1
LEG 37 BASALTS
In an important recent paper, Blanchard et al. (1976) have published analyses of basalts
from three DSDP sites ranging from 190-35 km away from the ridge axis in the FAMOUS
area (approximately 16.5-3.5 Myr old). Most samples came from the 35 km site, but all
samples were broadly similar and important age-related systematic changes in lava chemistry
(other than low-temperature alteration) have not been reported. Blanchard et al. used cluster
analysis to group their samples into five distinct groups, one of which they considered
parental to two of the other groups. For convenience the chief characteristics of the various
groups are listed below:
Table 1.
Approximate chemical content
Group Nature
I
Remarks
TiO,
(per cent)
A1203
(per cent)
MgO
(per cent)
14.2-17.2
6.0-9.3
0.7-1.2
18.0-21.6
9.8-14.7
6.5-8.5
11.5-24.3
0.35-0.53
0.4-0.55
IV
Plagioclase t Olivine
Phyric
Plagioclase Accumulative
Picrite (olivine
accumulative)
Olivine Phyric
14.6-15.3
9.2-11.3
0.6-0.74
V
Plagioclase Phyric
15 .O- 15.5
7.7- 8.6
0.7-0.9
I1
111
Frequent 'mantle' 01.
xenocrysts
Some xenocrysts
01. xenocrysts and
phenocrysts
01. phenocrysts
invariably present
Sometimes aphyric
640
4.2
E. G. Nisbet and C. M. R . Fowler
PETROGRAPHY
Several important conclusions may be deduced from the petrographic descriptions of these
rocks. Type 1 basalts frequently had bimodal textures, with a textural and compositional
hiatus between cores of large glomerocrysts (crystal aggregates) of plagioclase (Ang1 - g 5 )
and crystal rims and the fine-grained groundmass (containing plagioclase An40-50). This
compositional hiatus indicates a probable temperature drop of the order of 30°C or more
during multistage crystallization of the plagioclase (Humphries 1972). The most likely
explanation of this is that the glomerocrysts grew before eruption, perhaps in some sort of
magma chamber or conduit, while the groundmass crystallized on eruption. In addition to
plagioclase the groundmass contained clinopyroxene and minor olivine.
Type 1 basalts also contained irregular corroded olivine crystals with prominent kinkbands. These olivines were recognized as xenocrystal (foreign to the rock) by Blanchard
et al. (1976) because of these features, which indicate that they were out of equilibrium
with the melt and that they came from a deformed source, This source may have been the
mantle, though it should be noted that deformation of cumulus olivine crystals at the base
of a large magma chamber can also occur (Cameron, private communication; Jackson 1961,
p. 20). Blanchard et al. also figured a cumulus gabbro xenolith (rock fragment) from their
Type 1 basalts. Such a xenolith must have formed in the crust, not the mantle, and must
have been entrained by the rising liquid.
Type I1 basalts contained large amounts of accumulated plagioclase phenocrysts, as well as
xenocrysts of deformed olivine and gabbro fragments in some specimens. It is probable that
plagioclase floats in an oceanic basalt liquid (Bryan & Moore 1977), implying that Type I1
basalts were tapped from the top of a magma chamber, or that they represent some process
of concentration of lighter crystals. The groundmass was fine-grained, implying rapid final
chilling.
Type 111 basalts were characterized by abundant xenocrysts of olivine (derived from
deeper levels) and phenocrysts of olivine (i.e. olivine which crystallized from the liquid).
The groundmass of the rocks includes glass, implying rapid final chilling.
Type IV basalts have been divided into high, intermediate and low TiO, subtypes by
Blanchard et al. The lower TiOz samples were regarded by Blanchard et al. as probable
parents to the Type I1 and Type 111 lavas, and similar to the parent liquids of Type I basalts.
The rocks contained some xenocrysts of what may be mantle-derived olivine, but these were
not abundant. The low TiOz subtype contained equant phenocrysts of olivine in a largely
crystalline groundmass of plagioclase and clinopyroxene, implying crystallization at shallow
levels rather than in some magma chamber, or in the cores of pillow lavas. Some of the low
TiO, rocks were transitional to Type 111 picrites.
Type V plagioclase phyric basalts contained early crystallizing plagioclase with minor
olivine. The textures suggest that in the groundmass clinopyroxene, plagioclase and perhaps
olivine crystallized out together, implying shallow levels of final cooling (see later).
4.3
DEDUCTIONS FROM PETROGRAPHY
The presence of possible mantle xenocrysts in the Leg 37 basalts suggests that parent liquids
were mantle-derived, and were capable of entraining many mantle crystals as they rose.
Orthopyroxene xenocrysts have not been reported; this may be because they probably
react rapidly with the melt on ascent, or because orthopyroxene is not a liquidus phase in
the source. The implication of the xenocrysts is that important partial melting took place
in the mantle, though not to the extent of forming a 100 per cent molten ‘chamber’ (which
would lead to a xenocryst-free liquid unless the crystals were derived from the country rock
Mid-Atlantic Ridge at 37and 45" N
64 1
on ascent). The cumulus gabbro xenoliths (rock fragments) imply that the injection of the
liquids may have been through a roof of pre-existing high seismic velocity (c. 7-7.5 h / s )
gabbro, as the gabbro was probably broken off by the ascending liquid. Thus the most
likely site of final segregation and rapid ascent of the liquid is near or below the mantlecrust boundary or in layer 3B (Fig. 1).
The xenocrysts (about one-third of the rocks described by Blanchard et aZ. contain xenocrysts or xenoliths) place an interesting constraint on the dimensions of any postulated
high-level magma chamber. The xenocrysts often consist of single olivine crystals, with or
without chrome spinel inclusions, up to 0.5 cm long; larger xenoliths also occur. Biggar &
Clarke (1976) have investigated the reaction rate of xenocrysts in synthetic basaltic melts.
Their observations imply that single crystals would not be expected to remain out of
chemical equilibrium with a basaltic melt for more than a few days. Even if this time is a
matter of years rather than days, any process which brought the crystals from the base of
the crust or deeper must have been relatively rapid. If a large crustal magma chamber
existed, residence times of the liquid in the magma chamber would be very long. Magma
chambers of the size proposed by Hekinian et al, (1976) and Bryan & Moore (1977) would
contain los to 5 x lo5 times the average annual liquid increment from below if melting is
a continuous process. Thus for xenocryst-bearing rocks at least, the existence of a large
magma chamber is very doubtful.
The typical presence of phenocrysts (as opposed to xenocrysts) does, however, imply
that considerable crystallization must have taken place during ascent and eruption of the
magma. Zoning in phenocrysts of Type I basalts implies that the lavas had a multistage
history including a period of slow cooling before eruption. The Type I1 accumulative plagioclase also supports the existence of a melt chamber. However, most basalts have a finegrained or glassy groundmass, implying either that they were erupted on the surface
(probably the dominant process) or that they were intruded into already cold material.
To summarize; a three-stage history may be deduced from the petrography:
(1) The existence of melt near or below the mantle-crust boundary.
(2) Separation of the liquid, and intrusion through pre-existing gabbroic crust to some
higher level, with some liquids spending a short time in a melt chamber or differentiating and
crystallizing during flow or intrusion.
( 3 ) Final rapid chilling, most probably by extrusion for the rocks studied.
4.4
M A J O R - E L E M E N T C H E M I S T R Y O F LEG 3 1 B A S A L T S
We shall now examine the chemistry of the rocks analysed by Blanchard et al. in terms of
the CMAS diagrams discussed above (Figs 2 and 3 ) .
4.4.Z Type I V and Type IIZ basalts
Type 1V basalts were thought to be c m e to parent liquids by Blanchard et al. Fig. 3 shows
that on all three projections the Type IV basalts lie about half-way between the 1 atm and
10 kb initial liquids generated by the partial melt of a four-phase mantle source. They
would most easily be produced by partial melting at around 5-8 kb. The diopside projection
and the enstatite projection illustrate this. The rocks lie on what appears to be an olivine
control line, liquids moving away from olivine in composition as they precipitate it. A parent
liquid produced by partial melting at, say, 5-8 kb would precipitate olivine if it rose rapidly
to the surface. Type I11 basalts (which contain abundant olivine crystals) show an extensive
olivine control line clearly seen in the diopside projection.
22
642
20
30
MS +
+CS
Fig. 3 (a)
30
LO
M--+
tC3A
Fig. 3(b)
Mid-Atlantic Ridge at 37 and 45" N
643
2
A23'
1 otm
LO
-'ZS3
30
'0
M2S 4
Fig. 3(c)
Figure 3. Rocks analysed by Blanchard et ~ l (1976).
.
Projections as Fig. 2. (a) Projection from olivine.
Top figure: open circles Type I basalts, solid circles Type I1 basalts, crosses Type V basalts. Bottom
figure: open circles Type I11 basalts, solid circles Type IV basalts. (b) Projection from diopside. Top
figure: Types I , 11, V. Bottom figure: Types 111, IV. Symbols as for Fig. 3(a). (c) Projection from
enstatite. Top figure: Types I , I1,V. Bottom figure: Types 111, IV. Symbols as for Fig. 3(a).
The olivine projection shows much the same story. If Type IV and Type I11 basalts
crystallized at shallow level after rapid ascent, plagioclase would probably be the second
phase to crystallize after olivine, while orthopyroxene would probably follow olivine if the
liquid were trapped and cooled at depth. Blanchard et al. indicate that in Type IV basalts
plagioclase did follow olivine, indicating a high-level process. Crystallization at depth would
produce gabbros.
4.4.2 Types I , ZZ and V basalts
The three projections (Fig. 3) all show Type I basalts clustering about the 1 atm. invariant
point, and there can be little doubt that these basalts have fractionated at very high levels
E. G. Nisbet and C. M. R . Fowler
644
(sea-floor = about 0.25 kb). Most rocks fall in the olivine-plagioclase fields, and these
minerals would dominate basalt mineral assemblages. Many rocks have textures indicating
simultaneous crystallization of several minerals. The ‘overshoot’ into the plagioclase phase
field shown by the diopside projection is probably a consequence of the presence of plagioclase phenocrysts in these rocks, and perhaps a failure of the generalized phase boundaries
drawn.
Type I1 basalts fall mainly in the plagioclase phase field in both diopside and olivine projections (the enstatite projection is subject t o projection errors in these very plagioclase-rich
rocks). This almost certainly reflects the accumulation of plagioclase phenocrysts which
controls the composition of these rocks, and this accumulation may have occurred at high
level.
Type V basalts, like Type I basalts, fall close to the 1 atm. invariant point. It is possible
that some of these rocks may contain calcium-poor pyroxene in the groundmass, or that
they may have slightly atypical water contents.
4.5
S U M M A R Y O F M A J O R - E L E M E N T P H A S E R E L A T I O N S H I P S FROM LEG 3 1
BASALTS
A good case can be made out for the generation of some Type IV basalts as partial melts
of mantle material, with melting taking place perhaps at around 5-8 kb. These parental
melts could have suffered three fates:
(1) They could have been erupted direct to the surface and chilled sufficiently rapidly to
preserve the chemistry of 5-8 kb melting.
( 2 ) They could have been erupted rapidly to shallow level and then become trapped for long
enough for phenocrysts to grow (either in a small high-level magma chamber, in a thick
flow or even in a pillow core). If trapped as liquid at high level they would have precipitated
olivine rapidly. This may have produced cumulate-olivine enriched Type I11 lavas (also rich
in settled xenocrysts). The residual liquid may then have evolved towards the 1 atm.
plagioclase/olivine boundary (see diopside projection) and eventually to the 1 atm. invariant
point. It is also possible that upward flotation of plagioclase may have taken place, giving
Type I1 accumulative plagioclase rocks.
(3) Some of the liquid may have been trapped at the base of the crust or moved up slowly.
This would have led to precipitation of olivine, plagioclase, clino- (and ortho-) pyroxene as
the liquid evolved (the order depending on depth) to produce gabbro. Depending on the
shape and size of the liquid pool this gabbro would have cumulate to near-isotropic texture.
Blanchard et al. noted that Type I and Type V lavas could not be derived from a Type IV
parent. This implies that these lavas had a different magmatic history and could not have
shared a common magma chamber with the other types. Furthermore, Type IV basalts could
not have passed through a large magma chamber and retained the chemistry of 5-8 kb
melting. Thus it is unlikely that these rocks passed through a large magma chamber of the
sort proposed by Cann (1974) and others.
4.6
F A M O U S A R E A BASALTS
The above discussion has been exclusively on basalts from DSDP Leg 37. Two objections
can be raised to using the analysis above to deduce the present-day workings of the MAR:
(a) the Leg 37 basalts may not be typical of the basalts erupted today in the median valley,
and (b) only whole-rock analyses were considered. It is possible that these are not wholly
representative of liquid compositions.
645
Mid-Atlantic Ridge at 37and 45" N
/
"
/ "
-
"
Y
Y
Y
"
20
30
cs
(a)
\
/,
A
/
Y
"
"
2s
5
0
0
OPX
LO
MS +
Y
c
Y
.I
"
30
Figure 4. Glasses analysed by Bryan & Moore (1977). Points selected from data to demonstrate range of
composition present, including first sample from each dive and samples with high MgO content. All points
recalculated with 10 per cent of iron allocated as Fe,O,. Projections as Fig. 2. Triangle: average liquid
inclusion from Donaldson & Brown (1977). Cross: sample 528 before recalculation of iron. Allocation of
20 per cent of iron as Fe,O, would move the trend farther to the left in Fig. 4(b). (a) Projection from
olivine. (b) Projection from diopside. (c) Projection from enstatite.
646
E. G. Nisbet and C. M.R. Fowler
To study these problems we have used glass analyses presented by Bryan & Moore (1977).
These authors have studied fresh rocks collected from the floor of the present-day median
valley in the FAMOUS area. The rocks thus represent the top of the lava pile - i.e. the most
recent products in the axis of the valley, and slightly older lavas (but younger than the
underlying crust) away from the axis. Bryan & Moore investigated fractionation relationships by analysing basaltic glasses by electron microprobe. This offers a very powerful
method of directly measuring the compositions of basaltic liquids. The drawbacks of
electron-probe glass analysis are well known: often the sample deteriorates under the probe
beam; because of this and for statistical reasons light elements (e.g. Na) cannot be easily
measured; and FeO/Fe203 ratios cannot be directly measured. Nevertheless, Bryan & Moore
(1977) demonstrated that their analytical precision and accuracy were good (except perhaps
for Na, Mg and Si). We have recalculated their FeO analyses, allocating 10 per cent of the
iron on an atomic basis as F e 2 0 3(Hekinian et al. 1976, p. 104). Fig. 4 demonstrates the shift
this recalculation causes on one point in the CMAS projection. For all points this shift is
small.
Fig. 4(a, b , c) demonstrate the controls on evolution of the liquids analysed by Bryan
& Moore. The diopside and enstatite projections imply a strong olivine control on liquid
evolution (probably with plagioclase). If some of the liquids represent primitive melts, then
liquid 528 would represent a basaltic liquid generated by partial melting of peridotite at
5-10 kb; the other liquids can be derived from liquids similar to 528 by fractionation. The
olivine projection shows what is probably control by both plagioclase and clinopyroxene
(or perhaps by resorbed xenocrysts of orthopyroxene), suggesting that olivine, plagioclase
and clinopyroxene were all important in the evolution of the glasses, and supporting Bryan
& Moore’s conclusions.
The simplest conclusion that may be drawn from these basaltic glasses is that the liquids
were originally generated at pressures around, say, 5-8 kb by partial melting of a mantle
source, and then rose to the surface fractionating olivine, and crystallizing plagioclase and
clinopyroxenes on ascent, to give the high-level products seen. This interpretation is entirely
compatible with our interpretation of the Leg 37 rocks.
Donaldson & Brown (1977) have suggested that a liquid similar in composition to an
Archaean basaltic komatiite (i.e. rich in MgO and CaO) could be parental to some oceanic
tholeiites. This interesting possibility is not supported by Bryan & Moore’s data (see Fig. 4),
but may apply elsewhere in the MAR.
An important caveat must be introduced to the argument here. In the FAMOUS area, as in
the Leg 37 rocks, xenocrystal material is widely recognized (e.g. Hekinian et al. 1976). It is
probable that ascending melts carry with them a large burden of xenocrystal material. If
significant resorbtion of these crystals took place and they contributed to the chemistry
of the ascending liquid, then conclusions based on equilibrium partial melting and
crystallization assumptions will be somewhat in error.
4.7
T R A C E - E L E M E N T G E O C H E M I S T R Y O F F A M O U S A N D LEG 3 1 R O C K S ; R E S U L T S
O F GEOCHEMICAL MODELLING
The trace-element geochemistry of rocks from the MAR has been discussed by many
authors, notably Kay er al. (1970); Schilling (1975); O’Hara (1977); O’Nions & Pankhurst
(1974); O’Nions, Pankhurst & Gronvold (19?6), Yoder (1976) and in important recent
papers by Pankhurst (1977), Langmuir er al. (1977) and Flower er ~ l(1977).
.
The reader is
referred to these authors for extensive discussion of the problem.
Langmuir er al. found that rare-earth element (REE) patterns from fresh basaltic glasses
collected in the FAMOUS area were very variable, often enriched in light REE. They con-
Mid-Atlantic Ridge at 3 7 and 45" N
647
cluded that the large variations in trace-element abundances and in La/Sm and La/Yb
ratios suggested that the least fractionated FAMOUS basalts could have been extruded
directly from the source of melting without residence in a magma chamber. Their work
demonstrated that successive eruptions in one small area of the rift valley could show wide
variations in trace-element chemistry over a short span of time; this precluded the derivation
of basalt liquids from a single magma chamber. White & Bryan (1977) noted the great
ratios in analysed basalts from FAMOUS, implying a homogeneous
uniformity of "Sr/''Sr
source mantle, and also that not all REE variation in their samples could be accounted for
by closed-system fractional crystallization - other processes such as variation in extent of
partial melting must play a part.
Blanchard et al. (1976) and Puchelt & Emmerman (1977) studied REE in Leg 37 samples.
Blanchard et al. showed that Leg 37 basalts,like those from FAMOUS, showed considerable
variation. They concluded that several different parental magmas must have existed to produce the variation shown. Type I1 and Type 111 lavas (see above) could be derived from
Type IV, but Types I and V basalts could not be related to a common parent or to each
other.
Langmuir et al. and Pankhurst (1977) have investigated the relative effects on the REE
patterns of basaltic liquids of processes such as batch partial melting, incremental melting
(removal of melt increments as they form), zone refining (the process of melt rise by assimilating material above it and simultaneously depositing cumulates behind it), fractional
crystallization and magma mixing on the REE patterns of basaltic liquids. Langmuir et al.
concluded that their analysed basalts were derived from a suite of liquids produced by
varying degrees of partial melting of a rising mantle source with continuous but incomplete
removal of melt as melting proceeded, as well as varying extents of batch partial melting
and zone refining. It is difficult to quantify the degree of partial melt needed to produce
the observed REE patterns, but they demonstrated that for three FAMOUS samples 825 per cent partial melt of a source nearly twice chondritic in REE composition would yield
REE patterns similar to those observed. Bence & Taylor (1977) estimate 5-20 per cent
partial melting in Leg 37 rocks, using a similar source. Recent work on Archaean komatiites
(Bickle el al. 1976) has suggested that part of the mantle (albeit a very long time ago)
had a nearly chondritic REE content ; if the source mantle in the MAR is similar this would
imply rather low degrees of partial melt.
Some basic conclusions may be drawn from the controversial trace element work:
(1) Recent studies indicate that the basaltic liquids could not all be derived from a single
magma chamber. They must have been able to ascend separately to the surface from
different parents. Some may have been directly derived from the melting source.
(2) It is difficult to gauge the degree of partial melt needed to produce the liquids, but it
may be in the region of 5-25 per cent.
( 3 ) High-level fractionation is a very important process in influencing the composition of
many of the erupted liquids, but to produce the compositional range observed variable
partial-melting processes must also be invoked.
4.8
BASALTS F R O M 45" N
Basalts from 45" N have been studied in very much less detail than those from the FAMOUS
area. However, the seismic structures of the two areas are similar (Fowler 1978) and we will
thus briefly examine the rock types to see if they too are similar.
Muir, Tilley & Scoon (1964) described a large collection of rocks dredged from the MAR
median valley at 45" N . All were olivine basalts or dolerites; they were typically olivine-
648
E. G. Nisbet and C. M.R . Fowler
tholeiites and chemically very similar to those from the FAMOUS area (A1203
c . 15 per cent,
MgO 8-1 0 per cent approximately). Interestingly, some carried xenocrysts of highly
magnesian olivine, highly calcic plagioclase and chrome spinel. These were found in about
one-fifth of the basalts, and in the olivine-enriched dolerites. Plagioclase xenocrysts displayed marked resorbtion, and both olivine and spinel exhibited very strong marginal zoning.
The minerals characteristically occurred as large isolated single crystals or as small monomineralic aggregates of olivine or plagioclase. Olivine xenocrysts were rich in Ni and had
compositions around F090-91. Many crystals showed well-developed glide lamellae on [ 1001
and carried large inclusions of deep brown chrome spinel. Plagioclase xenocrysts ranged from
An87-92. No xenocrystal pyroxenes were found. Rarely, small xenoliths (rock fragments, as
opposed to crystals, strange to the melt) were found.
The chemistry of the rocks has been discussed by O'Hara (1968). In the CMAS projections they fall in similar fields to the FAMOUS area basalts, and there is np reason to
suppose that an extensive collection from the 45"N area would not produce exactly the
same diversity as in the FAMOUS area. Nicholls & Islam (1971) have discussed the FEE
patterns of rocks from 45" N - again they are similar to the FAMOUS area results and
display typically flat normalized curves.
Thus the available information demonstrates that the basalts erupted at 45" N are
probably chemically closely similar to those found in the FAMOUS area; it is notable that
both sets of rocks contain xenocrysts of possible mantle-derived olivine and plagioclase of
deep-level origin. In the following discussion we shall assume that processes at 45 and 37" N
are broadly similar and discuss the two areas together.
4.9
SUMMARY O F PETROLOGICAL CONSTRAINTS
(1) Petrography : the frequent xenocrysts of deep-level material probably mean that many
of the rocks only spent a very short time in a high-level magma chamber. The occasional
xenoliths further suggest that it was physically possible to bring up fairly large pieces of
deep-level material directly to the surface. The presence of cumulus gabbro xenoliths
suggests that gabbro is present under the median valley.
(2) Experimental petrology indicates that the simplest interpretation of the rock chemistry
is that partial melting takes place at a depth of 5-8 kb; however, a wide variety of alternative suggestions may be made, ranging from initial melting of garnet lherzolite at great
depth, followed by segregation and rise of the picritic (Mgrich) liquid from c. 30 kb; to
extensive melting of pyrolite at intermediate depth (1 0-20 kb).
(3) Trace-element analysis shows that much of the chemical variation seen can be ascribed
to high-level fractionation; however, some variation is too great to be explained by this,
implying that separate magma batches can reach the surface without mixing, and that
partial melting controls much of the REE patterns seen.
5 Models of ridge processes; melting and crystallization
We shall now consider the available models of ridge processes using these geophysical and
petrological constraints. High-level processes will be discussed first, since they are better
constrained, then we will discuss the problems of melting.
The most popular model of crustal processes at mid-ocean ridges is the 'infinite onion'
model (Cann 1968, 1970, 1974). This model has been adopted by many authors (e.g.
Bryan & Moore 1977; Hekinian et al. 1976; Rosendahl 1976). An alternative (Fowler 1976;
Nisbet 1974) is that on slow-spreading ridges there is no large magma chamber and that melt
batches rise separately to the surface. We shall call this the 'infinite leek' model.
Mid-Atlantic Ridge at 37and 45” N
649
5.1 T H E I N F I N I T E O N I O N
Cann (1970, 1974) proposed that a high-level (crustal) magma chamber exists in the typical
mid-ocean ridge; he termed this the ‘infinite onion’ as it continuously peels off layers of
crust at the edges while it is fed with liquid from below. Conceptually the model is very
simple: as the ridge spreads successive increments of lava are erupted to form a roof to the
chamber and simultaneously crystals are plastered to the walls of the chamber. The effect is
to produce a three-layer crust - a top of basaltic pillow lavas and dykes, a middle unit of
isotropic gabbro formed by gradual freezing of the margins of the magma chamber, and a
lower unit of cumulate gabbro formed on the base of the chamber (Fig. 6a).
Cann’s model is very attractive; it is elegant and it provides a suitable explanation of the
apparent seismic layering of the crust. It has also received powerful support from recent
observations on the East Pacific Rise (Rosendahl et al. 1976; Orcutt et al. 1975). Recent
work in the FAMOUS area has also been interpreted as supporting Cann’s model. Bryan &
Moore (1 977) and Hekinian et al. (1976) both noted the strong petrographic zonation of the
inner floor of the median valley. Regular compositional variations were observed from the
centre of the valley (which contained lavas rich in olivine phenocrysts, which were erupted
at relatively high temperatures and which had a ‘primitive’ chemistry) to the margins of the
valley (which contained lavas rich in plagioclase and clinopyroxene phenocrysts, which were
erupted at relatively cool temperatures and which had an ‘evolved’ chemistry). They interpreted this strong zonation as evidence for a large zoned magma chamber extending across
the width of the inner floor of the median valley (6 km - Bryan & Moore, Fig. 16). They
proposed that lateral differentiation within this large magma chamber could account for the
strong zonation of the erupted products on the floor of the valley; eruptions away from the
axis would tap a relatively differentiated and cool source in comparison t o axial eruptions.
They regarded the magma chamber as a steady-state feature, although it may occasionally
become blocked by settling crystals. All the surface lavas observed in the FAMOUS area
were much younger than the inferred spreading age of the crust on which they lay - this
would be consistent with the consequences of the magma-chamber model; but flank lavas
were rather older than central lavas, which is not explained by the magma-chamber model,
except by the argument that flank eruption is perhaps less common than axial eruption and
thus samples are likely to be older.
Ballard & van Andel (1 977) examined the topography of the median valley and noted
that subsidence after volcano construction in the centre of the median valley was greater
than could be accounted for by isostatic compensation. They interpreted this as evidence
for the collapse of a depleted magma reservoir, and in general supported Bryan & Moore’s
concept of a large magma chamber, occasionally becoming depleted. However, it could be
argued that this topographic evidence is better explained by a small high-level chamber
becoming exhausted, rather than a chamber the size of the whole median valley.
Thus there is considerable evidence which has been interpreted as supporting Cann’s
hypothesis in the MAR, and this is certainly at present the most popular theory for explaining the generation of new crust in the MAR.
There are, however, several very strong objections to the infinite onion as a process on
the MAR (though not on the EPR). They are:
(1) The seismic evidence, particularly from S waves. No magma chamber larger than about
2 km across at the most can at present exist beneath the MAR in the two areas studied.
( 2 ) The thermal constraints on the existence of a chamber make a large chamber very
unlikely. At the most it can only be about 1-2 km across in the FAMOUS area and at
45” N , and is probably smaller than this, if it exists.
650
E. G. Nisbet and C.M.R . Fowler
( 3 ) The presence of mantle xenocrysts in the erupted lavas implies that any lava batch only
spends a very short time in the chamber after its last contact with the mantle; a chamber
of the size usually suggested would mean that the average 'yearly' melt contribution would
spend 100000 to 500000 yr in the chamber, enough to destroy any xenocrysts.
(4) It is difficult to imagine how a large magma chamber could periodically allow apparently
parental liquids to pass through and yet maintain the chemistry of 5-8 kb melting equilibria. Similarly, it is probable that some of the lavas display geochemical variations that can
only be explained by variation in partial melt or by source inhomogeneity, not by any
process of open or closed-system fractionation; it is difficult to imagine how a lava can pass
through a large magma chamber without mixing of REE.
(5) The Rayleigh number for a magma chamber of dimensions between 1-6 km filled with
lava of viscosity 100 poise (about 30 stokes) is of the order of 10" to loz3,depending on
the size of the chamber and the temperature gradient across it. Even if the chamber were
filled with partial melt of viscosity 10" poise, the Rayleigh number would probably still be
in excess of 10". As the critical Rayleigh number is of the order of lo3,the chamber would
convect very vigorously indeed. It is therefore excessively unlikely that a large, quiet, layered
chamber could develop, sufficiently zoned to explain the strong zonation of the medianvalley lavas in both temperature of eruption and chemistry. This zonation is perhaps regarded as the strongest evidence in favour of a large magma chamber; convection considerations invalidate this. It could be argued that vigorous convection could easily mix mantle
xenocrysts into the melt (objection 3 ) ; however, the floor of the chamber would rapidly
re-equilibrate to the liquid under these conditions, removing the source of the xenocrysts.
Thus either the infinite onion does not exist under the MAR in the two areas studied, or
it does exist but is temporarily absent in both areas, or some other hypothesis must be
found. Any one of the objections outlined above could be regarded as fatal to the infiniteonion model; together they make a very strong case indeed. It must be noted though that
these objections only apply to the MAR, not to the East Pacific Rise.
5.2
THE INFINITE LEEK
An alternative model of the formation of the crust on the MAR may be constructed on the
assumption that no magma chamber exists, or if it does exist it is either narrow or transient.
This would be compatible with the S-wave results and with the thermal constraints. The
model is based (Fig. 5 ) on the discussion of Weertman (1971a,b) on the propagation of
magma-filled cracks in an elastic medium (the crust). Weertman demonstrated mathematically that if a tensile stress acts across a ridge axis and the crust is solid enough to behave as
an elastic plate, then a crack filled with melt will nucleate at the bottom surface of the
la1
lb)
Elastic Plate
Liquld Pool
t
/
\z\rf
-+
It1
Id)
-L7
7
1
7
Liquid Pool
Figure 5. Propagation of a liquid-filled crack (modified after Weertman 1971a). (a) Development of
liquid pool. (b)-(d) Initiation and propagation of crack. Note: the horizontal scale is greatly exaggerated
for clarity. A real crack would be very long and thin.
65 1
plate. This crack will increase in length and volume until it pinches its lower end shut and
rises as a ‘packet’ of melt plus entrained crystals towards the top surface. Eventually the
crack will stop, and the lava in it will freeze and become part of the new crust. Lava will be
extruded if a large enough crack reaches the upper surface, or if several cracks collect
together at the top. Intrusion of each crack will of course tend to happen most easily where
the crust is weakest at the site of the previous intrusion - i.e. at the spreading axis in the
centre of the median valley. The tendency of cracks, and thus ‘packets’ of melt, to stop
and collect at the top surface will produce a small very high-level magma chamber, comparable in volume to one or more cracks.
Mid-Atlantic Ridge at 37and 4 5 ” N
(a)
krn
6
krn
lo
(bl
1
6
Mantle
1
2
km
0
2
I
I
I
I
I
I
6
Ternporory hlgh- level
chambw
KEY
2
I
km
Melt
&
Trapped meit
Relict cryslais
6
72
8
0
Gabbros
I
10
Figure 6 . (a) ‘Infinite onion’ model of MAR in FAMOUS area, after Bryan & Moore (1977). The significant factor is the large magma chamber with cumulate base. (b) ‘Infinite leek’ model of MAR in FAMOUS
area, constructed from available seismic data (Fig. 1). Rising melt is probably trapped in the 7.2 km/s
layer, at the base of the elastic crust; from here it may inject to the surface leaving a cumulate residue,
or may become trapped at depth in small pockets crystallizing out as gabbro. Numbers: P-wavevelocity
in km/s.
652
E. G. Nisbet and C. M.R . Fowler
The 'infinite leek' model follows from this; it should be regarded as the limiting case of
ridge operation when the magma chamber is vanishingly small. Most mid-ocean ridges
probably fall somewhere between the infinite onion and infinite leek models - we would
consider the MAR to be closer to the latter. The model is as follows:
(1) Extensive partial melting begins in rising mantle at a depth of about 60 km (Forsyth
1977), and the degree of melting increases as the mantle rises. During this stage the melt is
in equilibrium with the parent rock.
( 2 ) Rising melt segregates from the mantle at depths of about 15-25 km (5-8 kb) and
travels rapidly and adiabatically upward as a diapir, probably precipitating mainly olivine.
(3) The melt reaches the base of the pre-existing crust at about 6-8 km below the surface.
At this stage it may be trapped by the overlying rigid mafic crust The pressure at this depth
is 1- 1.5 kb. The length of time the melt stays at this point depends on the quantity of melt
available, the tension in the crust and the strength of the crust (i.e. the time since the last
intrusion took place). If a small vertical magma chamber exists the liquid will mix into it,
and then eventually propagate to the surface by the method below.
(4) As the pool of liquid builds up a bottom crack can be nucleated into the basaltic plate.
This crack will most probably be in the same location as the last crack at the spreading axis.
The crack grows in length and the melt rises into it. As the magma enters the crack the base
of the crack eventually closes. Part or all of the magma will be absorbed into the crack, plus
disaggregated crystals of country rock and rare rock fragments if the process is very vigorous.
When the base of the crack closes it ascends rapidly into the basaltic crust to reach the top
surface, where a small volcano is constructed. If the volume of melt in the crack is large, or
if pre-existing melt from the last crack is present, the lava will spill out as a flow. It is
imagined that the volume of each crack is of the same order as an individual flow, or slightly
larger.
~
There are several consequences of this model:
(1) If a pool of melt is trapped under the crust it will begin to crystallize out as gabbro,
producing a body of gabbro the volume of the pool. Alternatively, if not all of the pool is
absorbed into the crack, a layer of cumulate crystals will be left behind at the base of the
crust, underplated onto it. Plagioclase may float and become attached to the roof of the
pool. Away from the axis the chance of melt becoming trapped at the base of the crust is
high, and thus a layer of gabbro would become plated onto the base of the crust. This may
explain the rapid thickening of layer 3 off axis (Fig. 1). At the axis the 7.6 km/s layer
probably consists of melt-rich mantle, away from the axis the 8.1 km/s layer is probably
cooler (and therefore denser) solid mantle, together with a top layer of cumulate peridotite.
The Moho may mark the level above which plagioclase becomes an important phase. Above
this will be the gabbro layer crystallized from trapped melt, containing both cumulate
gabbro (from the residue left by ascending melt) and isotropic gabbro (from trapped melt
pools too small to ascend). As gabbro accumulates and cools off-axis there will be a gradual
differentiation between melt and mantle.
(2) On the rapid final ascent of the crack, part of the liquid contents will be plastered to
the side walls; this will be preserved as a dyke and become part of the dyke screen.
(3) At the surface further fractionation may occur in a small magma chamber (within the
central volcano), and also during subsequent flow of the lava. The majority of lava flows
are probably relatively small (e.g. 500 m long appears to be typical (Bryan & Moore 1977;
Arcyana 1976); larger flows would show more differentiation and cooling at their tips.
F'icritic rocks would form at the base of the high-level chamber and also from the liquids
at the bottom of the crack.
Mid-Atlantic Ridge at 37and 45" N
653
(4) In a ridge spreading at 1.1 cm/yr (FAMOUS, 45"N), which is probably near the
minimum speed necessary to maintain thermally some sort of open-magma chamber at the
base of the crust, it is expected that at the axis cracks will intrude easily and a semipermanent weakness will be present. Thus it is likely that relatively primitive liquids will be
able to ascend easily directly from the mantle (at, say, 15 km) to the surface, without a long
delay at the base of the crust. It is also likely that during their last rapid ascent these liquids
may entrain many crystals of mantle or cumulative olivine and fragments of gabbro broken
off from the wall rock. Reaction between melt and crystals on ascent may somewhat alter
the chemistry of the source liquids.
(5) Rare-earth patterns would tend to be controlled mainly by the early partial-melting
processes, and by an approximation to closed-system crystal fractionation from the melt.
Thus we would regard the relatively flat patterns as being typical of the mantle segment
(which perhaps had a normalized Ce/Sm ratio slightly less than one), and variations in
Ce/Sm ratio as being probably mainly controlled by the extent of partial melting in each
magma 'packet', by the method of melt extraction and by fractionation at the base of the
crust and at high level.
(6) Ramberg & van Andel (1977) concluded that shifts of the spreading axis occur from
time to time, and that these shifts occur by the abandonment of one volcanic centre and the
initiation of another. This is consistent with the model suggested here: from time to time
it is to be expected that the central weakness will tend to close if there is an unusually long
gap between crack propagation events. If this happens a new set of cracks may propagate
elsewhere, probably up a fault weakness close to the old axis. A new volcanic edifice will
be created here, and act as a new locus of spreading.
(7) Some flank eruptions would be expected. It is probable that the faults which control
the walls of the median valley reach down through much of the crust. They must be a major
weakness in the crust and thus may be a favoured pathway for the ascent of cracks. Dyke
intrusion would also take place on many of the smaller faults, though many of these cracks
may not reach the surface.
(8) The strong petrographic and geochemical zonation of the inner floor of the median
valley noted by Hekinian et al. (1976) and by Bryan & Moore (1977) is easily explained,
by two separate mechanisms (Fig. 7). First, there is the observation that any lava outpouring
from the axis will tend to become more differentiated and cooler toward the tip of the flow
than at the source of the flow. Since the basement is spreading, the end of the flow is much
more likely to be preserved on the surface than the source, which will tend to be buried by
the next outpouring (Fig. 7). The tips of unusually large flows will have a much greater
chance of preservation on the surface, and thus it is to be expected that the axis of the
median valley will have relatively young 'hot' primitive lavas on the surface at any time,
Flank
I
Central volcann
i
Flank
I I ' I I I
A BC D C B A
I I I
I I I I
I I 1 I I I 1
Figure 7. Illustration of the inevitable petrogaphic zonation in the median valley that results from model
of Fig. 6(b). The first melt to reach the surface in any injection event is likely to be the most plagioclaserich, and will differentiate during flow, as do later products. Final eruptions before abandonment of the
central volcano are also likely to be o f relatively evolved liquid, as will be flank eruptions at the edge of
the median valley. Stippled: differentiated rocks. Unstippled: more 'primitive' rocks.
654
E. G. Nisbet and C. M. R . Fowler
while away from the axis the surface lavas will probably be rather older, 6cooler' and more
evolved. All surface lavas, however, will be much younger than the basement. Secondly,
liquids ascending from the mantle slightly away from the main axis of spreading would be
more likely to be trapped at the base of the crust. If eventually a large pool of liquid
accumulated and managed to escape to the surface, it would have a rather more evdved
chemical character than axial liquids and would probably be rather cooler. This mechanism
would also imply that mantle+derivedxenocrysts are much less common in flank eruptions
than in axial eruptions.
5.3
LOCUS O F PARTIAL MELTING
Solomon (1973) has shown that a zone of high S-wave attenuation under the Mid-Atlantic
Ridge extends to a depth of somewhere between 50-150 km, and in the Pacific Kausel
(1972, quoted by Forsyth 1977) showed that attenuation is best modelled by an attenuating
layer extending from beneath the crust to about 60-70km depth. Solomon & Julian
(1 974) modelled the temperature field and hence the P-wave velocity field in the upper
mantle near the axial zone. To account for the non-orthogonality of the nodal planes of
ridge-crest earthquakes they concluded that there must be a P-wave low-velocity zone in
the upper mantle 10-50 km beneath the MAR. Thus, from both these lines of evidence,
it is probable that the base of the zone of extensive partial melt is at about 20 kb depth.
The top of the zone is at a shallow depth beneath the crust (Fowler 1978; Steinmetz et al.
1977). The simplest conclusion of this and the petrological evidence is that partial melt
rises in equilibrium with mantle to depths of 15-25 km, after which it segregates and rises
rapidly through the mantle to the base of the crust (Fig. 8). This model is rather different
krn
20
30
10
0
10
20
30
GabSros/cumblates
lo
Monlle
&'
~
segregation of
'
partial m e l t ,
.
Rising p a r t i o i l y m o l l e n
mantle dtopirs
n
Iniliation of parfial
melhng
A
Rising m m t l e . unmolien
Figure 8. Model of melting processes under the ridge. Shape of zone of segregation of partial melt
estimated from petrological constraints and from thermal models (e.g. Sleep 1974, 1975). A: Initiation
of upward movement of mantle, from c. 200 km or more. B: Beginning of significant partial melting in
equilibrium with country rock. C : Segregation of melt. D: Formation of crust.
Mid-Atlantic Ridge at 3 7 and 45 ” N
655
from that of both Green & Liebermann (1976) and O’Hara (1968a), but is perhaps more
consistent with the available seismic evidence. The degree of partial melt is difficult to assess,
but REE evidence suggests that partial melts in the range 5-25 per cent would produce
the observed lava. If partial melting progresses to the extent that plagioclase is mostly
eliminated or has been eliminated, then a similar degree of melt is needed. Yoder (1976)
has discussed melting behaviour of garnet peridotite with temperature. If liquids rise
adiabatically to reach the surface at 1250”C, then final melting would have taken place at a
temperature not much greater than this. For sources similar to those discussed by Yoder this
would imply about 10-20 per cent melt. Of course other sources would have very different
melting behaviour.
Despite our preference for a shallow source, it is probable that either of the mechanisms
proposed by O’Hara (1968a) and Green (1971), Green & Liebermann (1976) will contribute
to the earlier stages of partial melting; it is possible that some magma batches may rise from
a greater depth.
5.4
COMPARISONS WITH FAST-SPREADING RIDGES; A N D THE EVIDENCE FOR
MANTLE ‘PLUMES’
The ‘infinite leek’ model may be regarded as the limiting case of ridge operation when the
magma chamber is vanishingly small; if the spreading rate increases a small magma chamber
would form at the base of the crust (layer 3B), and more frequent intrusions would
propagate from its top up the axis of the median valley. At high spreading rates the magma
chamber would be larger, and an ‘infinite onion’ would develop. There is considerable
seismic evidence that the East Pacific Rise contains a substantial crustal magma chamber
(RosendaN et al. 1976; Orcutt et al. 1975), and this is consistent with the available thermal
models (see above). Several major differences would thus be expected between the EPR and
the MAR:
(1) On the EPR a much more extensive cumulate layer would be developed than on the
MAR. The total thickness of cumulates may not be any greater than on the MAR, but the
cumulate layer would be much more regular and less patchy; it would be more easily
detected seismically. This is consistent with the suggestion of Moores & Jackson (1974)
that the thickness of layer 3B - the ‘cumulate’ layer - increases with spreading rate.
(2) Many more rocks produced by shallow and extensive crystal fractionation processes
would be expected on the surface in fast-spreading ridges. Clague & Bunch (1976) have
described ferrobasalts from the EPR which are consistent with this, and state that the EPR
is characterized by abundant strongly fractionated rocks.
(3) Incompatible element contents of average erupted rocks on the EPR would be expected
to be rather higher than in the FAMOUS and 4 5 ” N rocks, as the cumulation processes
in the magma chamber would partition these elements into the melt; this is supported by the
observation of Nisbet & Pearce (1973) that, on average, EPR basalts have much higher TiOz
content than MAR basalts, and that TiOz content of average oceanic basalts is correlated
with spreading rate,
(4) Fractionation and melting processes would be expected to be different in fast-spreading
ridges, and this would have important effects on REE patterns. At high levels on slowspreading ridges, with no open magma chamber, closed-system fractionation may dominate,
leading to minor changes in the Ce/Sm ratio (Pankhurst 1977). In contrast, on fast-spreading
ridges open-system fractionation would be expected to occur, with a very high ratio of the
volume of the chamber to the volume of an individual flow. This would lead to significant
effects on Ce/Sm ratios.
At deeper levels, on slowspreading ridges dynamic (Langmuir et al. 1977) melting
656
E. G. Nisbet and C. M.R. Fowler
processes may be expected to operate; in contrast on fast-spreading ridges incremental or
repeated partial melting may be important, as the ascent of magma may be rather easier
in a more rapidly flowing system. Incremental melting processes are capable of producing
liquids with lower Ce/Sm ratios than the source (Pankhurst 1977); thus a ‘depleted’ REE
pattern may result. The net REE pattern seen in erupted basalts is thus dependent on the
relative importance of four possible processes; it is therefore perhaps dangerous to speculate
on the presence of ‘plumes’ or on the homogeneity of the mantle given the poor understanding we have at present about the physical controls on melting and crystallization.
Although isotopic ratios are probably better indicators of source character than REE, it is
also possible that diffusion during inter-grain and grain-boundary movement in the early
stages of melting and disequilibrium melting (O’Nions & Pankhurst 1974) may alter initialsource ratios.
( 5 ) One final remark is that the pressure and temperature of initial melting would be
expected to be similar in both fast and slowly-spreading ridges. The ridge mantle ‘geotherm’
is probably mainly controlled by the rise of mantle from depth - in this respect it is similar
to ‘erosional’ geotherms seen in mountain belts on land. If the rising mantle starts from a
similar depth and temperature all over the world and rises nearly adiabatically it would
intersect the melting curve at the same depth irrespective of how fast it rises (assuming
conduction losses to be small). The other main control on the depth of the beginning of
melting is the shape of the melting curve, which is governed by chemical heterogeneity,
but probably does not vary too greatly.
Since the depth and temperature of melting is likely to be similar in both fast and slowlyspreading ridges, it follows that the extent of melting is also probably broadly similar. This
may be the reason for the lack of variation of the total thickness of oceanic crust with
spreading rate (Goslin et al. 1972), and for the general broad similarity of the Raitt-Hill
seismic layering throughout the world’s oceans.
5.5
COMPARISON WITH OPHIOLITE COMPLEXES
We have carefully constructed our model of ridge processes entirely from available oceanic
data to avoid the circular reasoning inherent in building up ‘ophiolitic’ models of the ocean
floor and then remarking on the surprising similarity of ophiolites to oceanic crust. Thus
we may use analogy with ophiolites as a test - though not a definitive test - of our model.
The literature on ophiolites is extensive and varied; we shall use the Troodos complex,
Cyprus and the Othris complex, Greece, for our test. Reference to the work of Moores &
Vine (1971), Smewing, Simonian & Gass (1975) and especially Menzies & Allen (1974)
and the important work of Allen (1975) shows that most of the features seen in these two
complexes can be duplicated in our model. Crude MgO balance calculations between basaltic
liquid, mantle source and peridotitic residue based on data in Menzies & Allen (1974)
suggest partial melts between 10-20 per cent or more for Troodos. Allen (1975) estimated
the size of magma pools as ranging from a few hundred metres to 2 km wide. Readers are
referred to the authors listed above for more extensive comparisons between our model and
ophiolite complexes; we believe that there are no fundamental disagreements.
6 Conclusions
The more important of our conclusions are summarized as follows:
(1) No crustal magma chamber larger than about 1-2 km wide can exist or needs to exist
in the MAR.
657
( 2 ) The petrographic and geochemical features of slowly-spreading ridges can be most easily
explained by a model involving periodic crack propagation into the crust from pools of
magma at the base of the crust (the infinite leek). On fast-spreading ridges a magma chamber
probably does exist (the infinite onion).
(3) Significant partial melting probably begins shallower than 60 km beneath the axis,
followed by liquid segregation at depths of 15-25 km. After this, liquid rises rapidly to the
base of the crust.
(4) Until more is known about the relative importance of incremental versus batch melting,
methods of melt extraction and open versus closed-system fractional crystallization, farreaching conclusions about mantle composition are not justified by REE evidence alone.
(5) Ascending liquids may entrain large amounts of crystals (most probably plagioclase or
olivine) from their source pools. Reaction on ascent may considerably change the chemistry
of the melt.
Mid-Atlantic Ridge at 37and 45"N
We believe that our model provides a satisfactory solution to the present conflict between
geophysical and petrological models of the MAR, and that it successfully explains all the
available petrographic and geochemical data from both 37 and 45" N.
Acknowledgments
We thank W.E. Cameron, K. G. Cox, D. H. Matthews, R. D. Beckinsale and V. Dietrich for
their very helpful criticism (not all of which has been followed!);K. Hsu, St. Miiller and
A. B. Thompson for providing research facilities; and especially the Royal Society, London
for European Exchange Fellowships. S. W. Richardson is thanked for allowing us to use his
petrological projection program, and C. R. Allen for allowing us to quote from his unpublished thesis.
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Mid-Atlantic Ridge at 3 7 and 45 a N
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