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Supplementary Information
Crustal thickness averages. The averages of the crustal thicknesses determined at
the harp and dixie locations, west of km 330, is 6.94 ± 0.12 km; crustal thickness
measured at abita, between km 330 and km 425, is 6.47 km; and the average crustal
thickness determined at the four locations east of km 425 is 5.6 ± 0.3 km. The ranges
indicated here are the standard deviation about the calculated means. Uncertainties of
the individual crustal thickness measurements are given in Fig. 3.
Line-1 Gravity Profile. We have constructed density models and calculated their
predicted gravity anomalies to test the consistency of our seismic velocity models with
observed gravity anomalies. We consider three basic models. In the first (Fig. S1), the
density structure of the mantle is based solely on temperature variation as determined
for a plate-cooling model with appropriate consideration of spreading-rate changes. The
density structure of the crust assumes a constant oceanic Layer-1/Layer-2 ratio and
crustal thickness as defined by the constraints from travel time modeling of the OBS
data along Line 1. This model is not in isostatic equilibrium, and it does not fit the
observed free-air gravity anomaly well. The second model (Fig. S2) employs the same
mantle density structure, but the crustal structure is imposed such that isostatic
equilibrium is maintained. This model fits the gravity data somewhat better then the
model of Fig. S1, but the fit remains poor. The model also does not satisfy the seismic
constraints on crustal thickness, with the crust in the east on average ~1.5 km thicker
than the seismically defined crust. The final models (Figs. S3 and S4) are motivated by
the mantle refraction results. We take the amount of low-density material needed to
attain isostatic equilibrium and distribute it within the uppermost mantle, reducing the
density there. For the model in the Fig. S3, we add this “gabbroic” component starting
with 0% at 30-km depth, increasing linearly upward toward the base of the crust. This
model yields a good fit to the gravity data, and it is consistent with the seismic
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observations on crustal thickness, with isostacy, with the requirement for a positive
velocity gradient in the upper mantle, and with our notion that this gradient is due to a
retained gabbroic component. A model where the distribution of melt begins at 60-km
depth (Fig. S4) fits the gravity somewhat better than the model of Fig. S3, but gabbroic
material is likely to transform to eclogite within the 30- to 60-km depth range, and so
the imposed crustal density of this added component is probably not physically
plausible. We indicate the root-mean-square misfit of the various models in Fig. S5.
In summary, we believe that the gravity results strengthen the notion that a
gabbroic component retained in the upper mantle provides a plausible explanation for
the large velocity gradients observed over the eastern portion of Line 1. Our preferred
density model, Fig. S3, is based solely on the crustal thickness observations and a
requirement of isostatic equilibrium. These requirements impart both positive velocity
gradients in the east and a transition from positive to approximately zero velocity
gradient near the location where the transition in mantle propagation properties is
observed from the mantle refraction data. This model fits the observed gravity data
better than an isostatically balanced lithosphere with uniform crustal thickness and
mantle density.
Conceptual model of melt retention. The spatial offset between the change in
mantle seismic propagation properties (km 300) and the location of the spreading rate
change from 13 mm/yr to 8 mm/yr is consistent with our conceptual model of melt
retention. This model, illustrated in Figure S6, is based on the widely held notion that
melt feeding crustal production at a mid-ocean ridge is drawn from the entire melting
region, which extends more than 100 km off of the mid-ocean ridge axis [e.g.
Spiegleman and Reynolds, 1999]. It is likely that a distributed, channelized melt
network evolves as melt extraction proceeds, perhaps resembling an inverted fractal tree
(Fig. S6a) [e.g. Spiegelman et al., 2001]. Disruption of this network due to the
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thickening of the thermal lid is most likely to affect the upper channels that drain the
off-axis portion of the network, trapping melt within that portion of the network (Fig.
S6b). This trapped melt eventually crystallizes into the lithosphere as it moves farther
off axis. In this scenario, the retention of off-axis melts, beginning at some point in
time associated with a decrease in spreading rate, would necessarily be manifest
underneath lithosphere of greater age than the spreading-rate change. We believe that
this offset, illustrated in Fig. S6c, provides the explanation for the observed seismicpropagation/spreading-rate change offset along FAIM Line 1.
An interesting corollary of this scenario is that the observed seismicpropagation/spreading-rate change offset provides some information on the width of the
melt zone. This width is predicted to be 1.0-1.5 times the depth of melting for passiveflow spreading models and ~0.5 times for buoyancy-driven flow models [Spiegelman
and Reynolds, 1999]. The observed offset of ~100 km therefore tends to favour
passively driven flow and suggests a melt-zone thickness of 60-100 km, consistent with
current thinking [Salters and Hart, 1989; Shen and Forsythe, 1995]. In addition, we
note that the preferential retention of off-axis melts, as illustrated in Fig. S6, is the basis
of our explanation for the widely observed correlation between thin oceanic crust and
increased concentrations of incompatible elements, as described in the text of the paper.
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Supplemental References
Salters, V. J. M., & Hart, S. R. The hafnium paradox and the role of garnet in the
source of mid-ocean ridge basalts, Nature, 342, 420-422, 1989.
Shen, Y., & Forsyth, D.W. Geochemical constraints on initial and final depths of
melting beneath mid-ocean ridges, J. Geophys. Res., 100, 2211-2237, 1995.
Spiegelman, M., Kelemen, P. B., & Aharonov, E. Causes and consequences of flow
organization during melt transport: The reaction infiltration instability in compactible
media, J. Geophys. Res., 106, 2061-2077, 2001.
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Supplemental Figure Captions
Figure S1. bottom: Density model based on seismic constraints on crustal
structure for crustal densities and plate cooling for mantle densities. Only the
upper 100 km of this 200-km-deep model is shown. This model is not in
isostatic equilibrium. The white line indicates the location of a smooth,
isostatically balanced crust. Note that the difference between the Moho based
on seismic constraints and the Moho of the isostatically balanced crust is ~1.5
km over the eastern portion of Line 1. top: Observed (black) and calculated
(blue) free-air gravity anomaly. The fit of the model to the data is poor.
Figure S2. bottom: As in Fig. S1, but now the model consists of an isostatically
balanced crust. The white line indicates the seismically constrained Moho. top:
The fit of the predicted response to the data is improved over the “seismic crust”
model, but it is still poor.
Figure S3. bottom: As in Fig. S1, but now the amount of low-density “crustal”
(or melt derived) material needed to isostatically balance the lithosphere is
distributed over the uppermost mantle, beginning with 0 added material at 30km depth and increasing the amount of added, low-density material linearly
toward the Moho. top: The fit of the predicted response to the data is good.
This model satisfies the seismic constraints on crustal thickness, it satisfies
isostacy, is satisfies the gravity observations, and it is consistent with positive
velocity gradients over the eastern portion of the model.
Figure S4. bottom: As in Fig. S3, but now the amount of low-density material
needed to isostatically balance the lithosphere is distributed from 60-km depth
toward the Moho. top: The fit of the predicted response to the data is
somewhat better than the “30-km” model, but a gabbroic phase below ~35-km
depth would begin a transformation to eclogite and so this density structure is
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difficult to reconcile with a simple addition of basaltic component due to
incomplete melt extraction.
Figure S5. bottom: Fits of the various density models to the observed anomaly.
top: RMS misfit for the various density models. Misfit for the “seismic crust”
model (Fig. S1) is not plotted, but is ~25 mGal.
Figure S6. Conceptual model of increased melt retention accompanying a step
decrease in spreading rate. a) Upwelling driven by separation of the lithosphere
(dark grey) occurs over a broad region, and the extent of the melt zone (red)
beneath an spreading centre is correspondingly large, 100 km or more. Melt
extraction occurs though a distributed, channelized melt network that focuses
melt from all portions of the melt zone to the ridge. Some melt is not
incorporated into the network (yellow), and this melt eventually freezes into the
lithosphere as gabbroic inclusions (blue). b) Thickening of the thermal lid due
to a decrease in spreading rate (at triangle) is most likely to affect the upper
network channels that drain the off-axis portion of the network, trapping melt in
that area. c) This trapped melt eventually crystallizes into the lithosphere as it
moves farther off axis. Retained melt is present underneath lithosphere of
greater age than the spreading-rate change, with the offset between the
occurrence of retained melt and the spreading rate change reflecting the halfwidth of the melt zone (horizontal line over triangle). This offset may provide
the explanation for the observed seismic-propagation/spreading-rate change
offset along FAIM Line 1. Also note in (c) that the melt which feeds the
spreading centre originates from upwelling primarily beneath the spreading
center.
Figure S7. Analysis of amplitude versus offset (AVO) pattern for Tecate.
black/dot) root-mean-square (RMS) amplitude for a 0.5-s window about the
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mantle-refraction phase. red) 13-point (12 km) average of the Tecate RMS
values. green) RMS amplitude within a 1-s window 1s prior to the first arriving
energy, i.e. RMS amplitude of the stacked background noise. The step
increase in noise at –325 km is related to a change from 3-fold stacking to 1fold. blue) predicted amplitude response for the continuous-gradient model
(Fig. 3). heavy black) predicted amplitude response for the velocity-step model
(Fig. 3). The velocity-step model predicts the amplitude decay beyond –250 km
better than the continuous gradient model. However, the magnitude of
anomalous (with respect to a 1D response) amplitude behaviour in the data
between –125 km to –225 km exceeds the magnitude of the difference between
the two predicted responses beyond –225 km, suggesting that inferences based
on AVO should be viewed with caution. Further analyses of the amplitude
response of the FAIM data are underway.
Figure S8. FAIM profile 420 for offset –90 km to 90 km. The amplitudes, A,
have been scaled by offset, X, so that Ascaled = A X.
Figure S9. FAIM profile cass for offset –210 km to 0 km. The amplitudes have
been scaled by offset.
Figure S10. FAIM profile Tecate for offset –400 km to -200 km. The amplitudes
have been scaled by offset.
Figure S11. Model parameters and fit statistics for the crustal thickness and
velocity measurements at each instrument-pair location. Seafloor and
basement depth were constrained with multi-beam and multi-channel seismic
data. A constant oceanic seismic Layer-2/Layer-3 ration was assumed and
Layer-2 velocity structure was adjusted to satisfy refraction traveltimes. For
each instrument, Layer-3 velocity and thickness were inverted for using a fourparameter model: velocity at the top and bottom of the layer, and Moho depth at
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distances of 100 km on either side of the instrument. For each instrument,
velocity at the top of Layer 3 was held fixed and the three remaining parameters
were inverted for. In each panel, black dots represent the results of individual
inversions for distinct values of the top Layer-3 velocity. These dots are plotted
as a function of crustal thickness beneath the instrument and RMS error
predicted for the inverted model. A well defined minimum with respect to crustal
thickness exists for each instrument. The average Layer-3 velocity of the best
fitting model is indicated (blue) for each instrument. This sensitivity analysis
forms the basis of the error bars shown for crustal thickness in Fig. 4.