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AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES
Published by AGU and the Geochemical Society
Article
Volume 1
June 1, 2000
Paper number 1999GC000028
ISSN: 1525-2027
A shallow, chemical origin for the Marquesas Swell
Marcia McNutt
Monterey Bay Aquarium Research Institute, 7700 Sandholdt Road, Moss Landing, California 95039
([email protected])
Alain Bonneville
Equipe Terre-OceÂan, Universite de la PolyneÂsie francËaise, B.P. 6570, Faaa AeÂroport, Tahiti, French Polynesia
([email protected])
[1] Abstract: Young hotspot volcanoes within plate interiors are frequently surrounded by smooth,
broad regions of shallow seafloor termed midplate swells. These swells are typically hundreds of
kilometers wide and can be more than a kilometer in elevation. The most frequently invoked
explanation for these swells is that they represent the thermal and dynamic surface uplift from rising
mantle plumes. Here we argue that buoyancy of a volcanic unit underplating the Marquesan
volcanoes just below the Moho, as imaged by a seismic refraction experiment, is capable of
producing much, if not all, of what has been previously interpreted as a thermal swell in both the
bathymetry and the geoid. The shallow compensation depth previously calculated for the swell based
on geoid observations is thus expected given that the compensation resides at 20-km depth. The
volcanic unit underplating the islands would also be expected to have a thermal effect, but its
contribution to swell uplift is never more than 10% of the chemical contribution. The predicted
increase in heat flow, on the other hand, is very large, but the thermal transient rapidly decays in the
first 2 million years after emplacement. If similar large volumes of underplated material underlie
other hotspot volcanoes, problems in explaining the lack of heat flow anomalies at later times and
negligible rates of swell subsidence can be alleviated. Although volcanic underplating is unlikely to
explain the entire swell signature for particularly large and wide swells, the results from the
Marquesas suggest that at least some portion of hotspot swells may not be thermal in origin. After
removing the geoid signature from the Marquesas volcanoes and swell, what remains in the geoid are
stripes of amplitude 1 m and wavelength 650 km trending normal to the fracture zone. The anomalies
are presumably relict from when this crust was created, but their amplitude is too large to be caused
by variations in isostatically compensated crustal thickness.
Keywords: Hotspot; marine geophysics; marine seismology; Marquesas Islands; geoid interpretation.
Index terms: Dynamics, convection currents and mantle plumes; marine geology and geophysics; regional and global
gravity anomalies and Earth structure.
Received October 29, 1999; Revised April 27, 2000; Accepted May 5, 2000;
Published June 1, 2000.
McNutt, M., and A. Bonneville, 2000. A shallow, chemical origin for the Marquesas Swell, Geochem. Geophys. Geosyst.,
vol. 1, Paper number 1999GC000028 [6929 words, 7 figures]. June 1, 2000.
Copyright 2000 by the American Geophysical Union
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1. Introduction
The Marquesas (Figure 1) represent one of
a number of young hotspot volcanic chains in
French Polynesia. Radiometric dates of rocks
collected from the islands [Duncan and
McDougall, 1974] and dredged from seamount flanks [Desonie et al., 1993] range
from only a few hundred thousand years for
a seamount southeast of Fatu Iva at the southeast end of the chain (R. Duncan, personal
communication, 1995) to 5.75 million years
for Eiao atoll at the northwestern end [Caroff
et al., 1999]. The islands erupted onto seafloor
50±65 million years in age, based on the
identification of magnetic lineations on the
flanks of the Marquesas swell [Kruse, 1988;
Munschy et al., 1996]. Therefore the Marquesas seem to be a classic hotspot chain like the
Hawaiian Islands, with the only difference
being the rather recent onset of volcanic
activity as compared with the long (>65 Ma)
record for Hawaii.
[2]
Like other hotspot chains on mature lithosphere, the Marquesas are surrounded by a
broader region of uplifted seafloor [Crough
and Jarrard, 1981]. The width of the Marquesas swell is difficult to estimate because its
flanks merge so gradually into the surrounding
seafloor, and the result can depend on the
choice of reference model for age correction,
which is still necessary for lithosphere of this
age (50±65 Ma). The most prominent part of
the rise is within ‹250 km of the center of the
volcanic chain [Fischer et al., 1986]. Sichoix et
al. [1998] estimate that the total width of the
swell is 1300 km based on a modal analysis of
depth anomalies.
[3]
The amplitude of the swell is also difficult
to precisely estimate on account of the large
amount of extrusive volcanism comprising the
Marquesan volcanoes on the crest. Sichoix et
al. [1998] estimate that the amplitude is
[4]
1999GC000028
1000± 1300 m from plotting the modes of
depth, but this technique does not entirely
avoid the problem of surficial crustal thickening by extrusive volcanism and sedimentation,
particularly near the center of the chain where
the bathymetry is plateau-like. Fischer et al.
[1986] used a more sophisticated filtering
technique to attempt to extract the swell
topography from the extrusive volcanism.
Their method relies on the fact that loads
placed on the surface cause the Moho to warp
down, whereas any swell topography that is
compensated by low-density material beneath
the Moho would cause the Moho to warp
upward. They estimate a swell amplitude of
1000 m, but their method begins to break
down if the compensation depth for the swell
approaches that of the Moho, which we will
argue soon is probably the case for this hotspot chain.
2. A More Precise Estimate of Swell
Amplitude
Probably the most precise estimate of the
swell topography was indirectly implied in the
analysis of Wolfe et al. [1994] that modeled the
moat stratigraphy along three multichannel
seismic cross sections of the chain (Figure 1).
When all of the surface topography was assumed to be a surface load on the elastic plate,
they obtained a depth for the flexural moat that
is 2.8 km too deep along the center part of the
chain, 2.3 km too deep along a more northern
transect, and 1.7 km too deep along a more
southern crossing. To bring the predicted moat
deflection into agreement with the observed
seismic reflection data, they assumed that a
buoyant load flexes the elastic plate upward
from below, thus negating the loading effect of
some of the surface relief. In actual fact, the
source for the buoyant uplift of the seafloor
could be any thermal or chemical alteration of
the crust or mantle beneath the moat that
renders the rock less dense than its surround-
[5]
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tu
Frac
138˚W
-2500
-1875
Depth (m)
-1250
re Z
one
12˚S
136˚W
-625
0
625
Figure 1. Bathymetry of the Marquesas Islands [Filmer et al., 1994]. Inset shows the location of the islands
in the central Pacific at the northern edge of French Polynesia. The red line (dashed) shows the expected
azimuth of the chain were it properly aligned in the direction of absolute motion of the Pacific plate with
respect to the hotspot reference frame. The green line (dotted) is the average trend of the Marquesas chain.
The white lines correspond to the location of the multichannel seismic lines from Wolfe et al. [1994]. The
central white line also lies along the seismic refraction survey of Caress et al. [1995] and is the cross section
in Figure 4. Thick black lines are identified magnetic lineations with ages in million years [Kruse, 1988;
Munschy et al., 1996]. Diamonds show heat flow stations [Stein and Abbott, 1991], with size of the symbol
indicative of the magnitude of the heat flow. The medium size symbols are heat flow values within 1 sigma of
the expected value for lithosphere of that age. The smallest symbols are heat flow less than 20 mW/m2. The
largest symbol is a heat flow value of 108 mW/m2.
ings. Wolfe et al. [1994] found, however, that
they could obtain all of the uplift necessary to
reconcile the model moat depth with what is
observed if they assumed that an anomalously
thick crustal root, as imaged in a seismic
refraction experiment [Caress et al., 1995] is
a buoyant load flexing the elastic plate up from
below. According to this scenario, the anomalous root is interpreted as volcanic underplating of the original crust by magma less
dense than the surrounding mantle but more
dense than the overlying crust that therefore
ponded at the Moho upon ascent. The inference has been that this underplating body
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was formed in the last 9 million years
during the current phase of Marquesas hotspot volcanism.
Although the amplitude of this relief supported from below was not specifically called
out in their analysis, we can calculate its
amplitude as follows. In the wavenumber domain the surface deformation Hs(kx, ky) caused
by the loading of an elastic plate from within or
beneath by a load U(kx, ky) is given by McNutt
[1983]
[6]
Hs (kx ; ky † ˆ (m
"
(2k†4 D ‡ ( u
u †
g
. U(k x ; ky †;
#
w †g
1
(1†
where u, w, and m are the densities of the
underplating load, water, and mantle, respectively, D is the flexural rigidity of the plate, k is
the modulus of the spatial wavenumber, and g
is the acceleration of gravity. In this particular
circumstance the first term in the bracketed
numerator, which takes into account the rigidity of the plate, is much smaller than the
second for wavelengths of interest (several
hundred kilometers), and the equation reduces
to the following simple expression in the
spatial domain:
hs (x; y† ˆ
( m
( u
u †
u…x; y†:
w †
(2†
The value for u(x,y) and the densities are
taken from Wolfe et al. [1994]. The amplitude
of the load was chosen to be consistent with the
seismic refraction data of Caress et al. [1995],
the density u (3100 kg/m3) was chosen to fit
the moat depth as imaged in the seismic reflection data, and the existence of three seismic
reflection crossings at the southern, central, and
northern end of the chain allowed the body to
be interpolated into plan view from the individual cross sections.
[7]
1999GC000028
Figure 2 shows the predicted surface uplift
hs and a plan view of the underplated load u
from Wolfe et al. [1994] that compensates it.
In this calculation, the densities of water and
mantle are 1000 kg/m3, and 3300 kg/m3,
respectively. The maximum height of the surface uplift above the depth of the surrounding
seafloor is 843 m. This value is quite consistent with the observation that failure to include
the buoyancy effect of loading from below
causes the moat to be 2.8 km too deep. That
is, if 843 m of surface topography is no longer
assumed to be a surface load and is removed
from the compensation calculation, the moat
isostatically rebounds 2.1 km using the densities assumed here. If then enough load is
placed below the moat to create a swell 843 m
high, the moat must shoal by the same
amount. The sum of these two effects is
therefore 2.9 km, in close agreement with
the finding of Wolfe et al. [1994] that the
moat is 2.8 km too shallow when all of the
surface topography is assumed to be a load on
top of the plate. The amplitude and the overall
shape of this feature are quite similar to the
parameters describing the Marquesas swell. In
particular, the underplating body is concentrated in a band of width 450 km centered
beneath the island chain and thus produces the
greatest surface uplift in that region. Crustal
thickening persists, however, to a total width
of 1000 km beneath the chain, but at much
reduced amplitude.
[8]
We consider this to be the best estimate to
date of the amplitude of the Marquesas swell,
at least within the region beneath the island
chain and its flexural moat sampled in the
seismic experiments. The base of the moat
sediments as imaged by seismic reflection
provides a less ambiguous reference horizon
than does the surface bathymetry, which has
been clearly paved over with volcanic flows
and sediments. The fact that our estimated
swell height is several hundred meters less
[9]
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1000
875
6˚S
750
625
8˚S
500 M
e
375 et
r
250 s
10˚S
125
0
12˚S
-125
-250
10000
6˚S
8000
6000
8˚S
M
e
4000 t
e
r
s
2000
10˚S
0
12˚S
144˚W
142˚W
140˚W
138˚W
136˚W
-2000
Figure 2. (top) Height of the Marquesas swell in meters and (bottom) plan view of the underplating that
provides its compensation. The distribution of the underplating is taken from Wolfe et al. [1994]. Their
underplating function has been shifted vertically downward by 1 km in order to provide a slightly better fit to
the seismic reflection and refraction data through the center of the Marquesas chain.
than previous estimates is in the right direction
given their likely biases introduced by volcanic topography. Note that this estimate of
swell height is not necessarily dependent on
the assumption made by Wolfe et al. [1994]
that the support for this surface relief is the
volcanic underplating imaged by Caress et al.
[1995]. If the swell were supported just beneath the crust, at the base of the lithosphere,
or deeper within the asthenosphere, the result
would be the same: about 850 m of surface
topography must be supported by buried, lowdensity material rather than by flexure of an
elastic plate loaded from above.
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In general, we can express the total bathymetry h (observed depth minus a best-fitting
plane in this region to remove the age effect
from cooling of the lithosphere) as
thermal origin, McNutt [1987] calculated that
the Marquesas plume would have had to
reheat the entire bottom half of the 50-Ma
lithosphere to asthenospheric values.
(3†
If much of the swell is instead compensated by the underplating crustal root (e.g., hcs
as opposed to hts), it would provide an alternative origin for the swell that would not
require such massive lithospheric reheating.
To test this hypothesis, we determine whether
the geoid signature from the underplating
body and its surface expression is capable of
producing the long-wavelength geoid anomaly
over the Marquesas islands. We consider the
geoid anomaly N to be the sum of the
following terms:
[10]
h ˆ hv ‡ hcs ‡ hts
in which hv is the change in elevation caused by
loading from above by the volcanoes and their
sediments, hcs is the chemical component of the
swell cause by lateral variations in mineralogy
(e.g., the underplating), and hts is the thermal
component of the swell cause by lateral variations in temperature, such as what might be
caused by lithospheric thinning or dynamically
maintained thermal anomalies in the asthenosphere.
The 843 m of uplift needed to match the
moat depth represents the sum hs ˆ hcs ‡ hts .
We need additional information to determine
what portion of the total the swell is indeed
compensated by the chemically differentiated
material just beneath the Moho, as assumed
by Wolfe et al. [1994], and what part of the
uplift arises from thermal anomalies deeper in
the mantle.
[11]
3. The Marquesas Geoid
The Earth's equipotential surface, the
geoid, rises 2 m over the swell, with much
larger local anomalies associated with the
volcanoes [Marsh et al., 1986]. On the basis
of the geoid to topography ratio, Crough
[1978] estimated the compensation depth of
the swell to be 30 ‹ 40 km. Using a Fourier
technique that accounts for the finite elastic
strength of the plate, Fischer et al. [1986]
concluded that the compensation could be as
deep as 45 km. Both of these estimates put the
compensation for the swell well above the
base of the lithosphere. In order to produce
both the height of the swell and the shallow
compensation depth via density reduction of
[12]
[13]
N ˆ Nv ‡ Ncs ‡ Nts
(4†
in which N is the observed geoid anomaly
obtained by subtracting a best-fitting plane
from the total geoid in the region of study;
Nv is the geoid effect due to the surface
volcanoes and their flexural compensation,
Ncs is the geoid anomaly from the chemical
portion of the swell and its compensation by
crustal underplating, and Nts is the geoid
anomaly from the thermal portion of the swell
and its compensation by deep processes associated with the hotspot, such as lithospheric
thinning or dynamically maintained thermal
anomalies in the asthenosphere. To the extent
that Nts is small or of such short wavelength
as to be caused by errors in the surface terms,
we would have to conclude that there is no
evidence in the geoid for other density effects
from the hotspot deeper within or below the
lithosphere.
[14] We begin by removing from the geoid
anomaly N (Figure 3 (bottom)) the geoid
effect of the volcanoes and their compensation by a flexed elastic plate. Let hv equal the
amplitude of the volcanic load in the bathy-
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e
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0
6˚S
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8˚S
10˚S
12˚S
144˚W
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140˚W
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Figure 3. (bottom) Geoid anomaly over the Marquesas Islands [Marsh et al., 1986]. (middle) Geoid effect
due to the surface loads (volcanoes and sediment minus swell from Figure 2) plus their compensation
assuming that the load flexes an elastic plate 18-km thick. (top) ``Swell'' geoid, obtained by subtracting the
geoid in the middle panel from that in the bottom panel.
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2
4
zs
water
1000 kg/m3
Depth (km)
surface load
ztc
swell
2650 kg/m3
6
8
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moat
Te = 18 km
z bc
10
flexed oceanic crust
12
14
underplating
16
mantle
3100 kg/m3
18
3300 kg/m3
zm
20
22
-400
-300
-200
-100
0
100
200
300
400
Distance (km)
Figure 4. Cross section through our model along the central seismic line in Figure 1. Only that portion of
the bathymetric load above the swell is supported by elastic flexure. The swell itself is compensated by the
underplating. The shape of the underplating is consistent with seismic refraction data, and the depth of the
moat agrees with seismic reflection data. The depths to density interfaces ztc, zbc, zs, and zm used in the geoid
calculation are indicated.
metry at average depth zv below the sea
surface which is elastically compensated by
flexure of the Moho at average depth zm.
Then in the wavenumber domain [McNutt
and Shure, 1986],
G( v w †
kg
Nv (kx ; ky † ˆ
.
exp( 2kz v †
. H v (kx ; ky †
ˆ
exp( 2kz m †
1 ‡ k4 (2†4 D
:
( m v †g
(5)
(6†
G is Newton's gravitational constant and
v is the density of the volcanic load (=2650
kg/m3). For Hv we subtract from the observed
bathymetry the swell displayed in Figure 2,
as it has already been shown that this component of the bathymetry cannot be flexing
the elastic plate from above based on moat
depth imaged in seismic reflection data
[Wolfe et al., 1994].
[15]
Figure 3 (middle) shows the geoid from the
volcanoes and their compensation, and Figure 3
[16]
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(top) shows the result of stripping this signal
from the observed geoid. This is the swell
geoid. There is a smooth, 2-m geoid signature
along the axis of the chain caused by the
underplating and its compensation, and perhaps
by deeper thermal and dynamic processes.
Some residual high-frequency signal is well
correlated with the position of the major volcanic edifices in the bathymetry. The correlation is
sometimes positive, sometimes negative. This
indicates that our bathymetric model is less than
perfect (including possible offsets between the
origin for the geoid grid and that of our bathymetry grid) and/or that our choice of a uniform
density and flexural rigidity (as required by the
Fourier approach) is not always valid.
Next we remove the geoid effect of the
underplating, which includes the positive geoid
caused by the surface uplift as well as the
negative geoid from the underplating, which
is more buoyant than the surrounding mantle.
Ideally, we would like to have an estimate of
hcs, but all we really know is the sum hcs ‡ hts .
However, if we assume that all 843 m of the
swell is chemical in origin and compensated
just below the Moho, we will obtain a lower
bound on the size of the swell geoid. To the
extent that any part of hs is compensated deeper
in the mantle, the amplitude of the predicted
swell would increase on account of the attenuation of the signal from the density deficit with
increasing depth of burial.
[17]
To a first approximation, when the wavelengths of the underplating load are long with
respect to the flexural wavelength such that the
rigidity of the plate can be ignored [Turcotte
and Schubert, 1982],
[18]
Ncs ˆ
2G
g
Z
z(z†dz;
(7†
where is the distribution of anomalous
density. Using the four main density interfaces,
1999GC000028
ztc, zbc, zs, and zm, representing the depth to the
top of the flexed crust off the swell, the base of
the crust before underplating, the top of the
swell, and the Moho at the base of the underplating unit (Figure 4), this integral becomes
Ncs ˆ
G
‰( m u †(z2m z2bc †
g
‡( v w †(z2tc z2s †Š:
(8†
If the entire swell is caused by the underplating
unit, the implied density of the underplating is
3100 kg/m3.
[19] Figure 5 shows the predicted geoid from
the swell and underplating, while Figure 6
shows a three-dimensional image of the ``deep
geoid'' that remains after subtracting this from
the swell geoid in Figure 3c. The underplating
geoid is spatially coincident with the underplated unit and reaches a maximum amplitude
of 1.9 m. Note that the deep geoid bears no
correlation with the 4000-m isobath, which
approximates the outline of the Marquesas
volcanic platform. The only structures left in
the deep geoid are the high-frequency anomalies in the area of the volcanoes caused by
errors in the bathmetry and density model, and
a much longer wavelength series of alternating
high and low stripes trending perpendicular to
the fracture zone.
[20] We believe that the geoid stripes are a real
signal for the following reasons. First, their
amplitude exceeds a meter, while the altimeters
used to measure geoid are accurate to better
than a centimeter. Second, all of the corrections
we have made to the data to remove known
effects (volcanoes, flexure, underplating) involve structures trending in the direction of
the Marquesas Islands. We cannot create a
fracture-zone-normal artifact. Finally, both of
the positive geoid structures lie mostly outside
the region affected by our modeling. All we
have done is revealed that the geoid low
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7
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6
5
8˚S
4
3
2
10˚S
M
e
t
e
r
s
1
0
12˚S
-1
Figure 5. Predicted geoid from the swell and the crustal underplating.
already apparent in the unprocessed geoid
(Figure 3) well north and south of the islands
is continuous as it passes beneath the islands.
[21] With respect to these stripes, most of the
Marquesas Islands have erupted into the geoid
low, although the southern end of the chain and
the Marquesas Fracture Zone Ridge [McNutt et
al., 1989] pass over onto the adjacent high to
the east. Had we assumed that some portion of
the 843-m-high swell were compensated even
deeper than the base of the Moho, as would be
the case if some of the topography were thermally compensated at the base of the lithosphere, the geoid effect in Figure 5 would be
even larger, to the point of producing a more
pronounced depression, but with the trend of
the Marquesas, in Figure 6. The placement of
the island chain with respect to the fracturezone-normal geoid stripes explains why the
Marquesas Islands appeared to be offset with
respect to their swell, with the geoid high being
well developed to the northeast and virtually
nonexistent to the southwest.
4. Discussion
[22] The foregoing analysis demonstrates that a
swell height of 850 m is required in order to
reconcile the depth of the Marquesas moat
with elastic loading models. The geoid height
over the swell suggests that the compensation
depth for the swell is 20±25 km. Wolfe et al.
[1994] have already shown that the mass of
Figure 6. (top) Map view of the final residual geoid obtained by subtracting the geoid effect in Figure 5
from the swell geoid in Figure 3 (top). The black lines and diamonds are the magnetic lineations and heat
flow stations from Figure 1. The trend of the geoid highs and lows remaining after making known corrections
for the Marquesas Islands and their chemical swell is similar to that of the magnetic lineations, suggesting
that they date from when the lithosphere was created at the mid-ocean ridge. The wavelength of the
undulations is 650 km. These geoid lineations have a distinctly different trend from either the absolute
motion of the Pacific plate (dashed red line) or the trend of the Marquesas Islands (dotted green line).
(bottom) Three-dimensional view of the same geoid.
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54
o
6S
59
52
o
8S
66
60
49
o
10 S
63
69
56
o
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o
o
o
142 W
o
140 W
o
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136 W
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,m
0
1
144 W
o
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o
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o
o
10 S
o
o
o
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o
o
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-2
W
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-1
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W
0
Geoid (m)
W
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136
1
W
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the underplating material imaged by seismic
refraction data at 22-km depth beneath the
Marquesas is sufficient to produce the required
shoaling of the moat if the average density of
the underplated material is so high that its
contrast with respect to that of the upper
mantle is only 200 kg/m3. A high density for
the root material is also supported by the high
seismic P-wave velocities observed in the root:
7.5±7.75 km/s where constrained by turning
rays [Caress et al., 1995]. These velocities are
more than 2 km/s greater than that for average
basalt and suggest that the density of the root
material greatly exceeds 2800 kg/m3.
[23] One possible explanation for the origin of
the underplated material beneath the Marquesan Moho is that it represents melt or some
combination of melt and normal mantle material that has ponded at the density discontinuity
between the crust and mantle. The melt would
be positively buoyant with respect to the mantle
but too dense to ascend through the crust. On
the basis of the petrology of glasses recovered
from the Marquesas Islands and the inferred
physical properties of the underplating, it is
possible that the underplating material is an
alkalic gabbro with an MgO content of 4%
(J. Natland, personal communication, 1999) or
some combination of mantle and gabbros with
higher MgO (D. Clague, personal communication, 1999).
[24] The high probability that the underplating
is magmatic in origin leads to the question of
what would be the thermal consequences of
this hot body on the elevation and heat flow
across the Marquesas swell. Figure 7 shows the
results of a simple calculation on the thermal
evolution of a hot magma body ponding at the
base of the crust at time zero. The thermal spike
is 3 km thick, on the assumption that the
underplating is some combination of cooler
mantle and melt sills. The temperature spike
decays extremely rapidly, on account of the fact
1999GC000028
that heat is conducted both towards the surface
and back down into the mantle. The heat flow
signature of the underplating (represented as a
fraction of the normal value Qo) peaks within
the first two million years, and falls to a very
low value soon after. The initial elevation effect
is rather small, less than 100 m, and also decays
rapidly. If we assume that the entire 7-km-thick
underplating body was initially molten, a very
extreme model, the heat flow spike reaches 5
times the background value and decays more
slowly, such that the heat flux would still be 2
times the normal value at 4 myr after emplacement. The maximum elevation effect is 300 m
and is still as large as 150 m at 4 myr after
emplacement.
The uncertainty in how the seismic velocities in the underplated body convert to density would certainly allow 150 m or even more
of the density effect of the underplating to be
thermal as opposed to chemical in origin.
Detailed Seabeam plots acquired on transits in
and out of the Marquesas Islands reveal a break
in slope at 20- to 50-m depth on several crossings. One possible interpretation of this slope
break is that it measures the amount of recent
subsidence in the chain. If so, the model with a
smaller melt fraction and less thermal subsidence would be preferred. However, in either
case (the mixed mantle and melt or the pure
melt lens), the majority of the buoyant support
for the swell is chemical as opposed to thermal.
[25]
What exists of heat flow data also tends to
support the model with a smaller melt fraction
to the underplated unit. The average of 10 heat
flow stations [Stein and Abbott, 1991] plotted
in Figure 1 is 53.3 ‹ 26.0 mW/m2. This average
matches the value for normal lithosphere of age
50 Ma [Stein and Stein, 1993], despite the fact
that the mean depth at those heat flow stations
is 700 m shallower than expected for normal
lithosphere of that age. The 1-sigma uncertainty
would allow the heat flow to be 1.5 times the
[26]
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1999GC000028
MCNUTT AND BONNEVILLE: MARQUESAS SWELL
0
1 my
Depth (km)
20
time 0
40
60
80
0
200
400 600 800 1000 1200 1400
Temperature (degrees C)
Q/Qo
4
3
2
1
0
5
10
15
20
Time (million years)
25
Swell Height (m)
100
50
0
0
5
10
15
Time (million years)
20
25
Figure 7. Model to calculate the thermal effect of the underplating assuming heat flows via conduction
only: (top) Evolution in the geotherm after a spike of hot material 3 km thick is injected at the base of the
Moho. Geotherms are plotted at time 0, 1, 2, and 3 million years. (middle) Change in heat flow with time,
expressed as the ratio of the normal heat flow just before the underplating event. (bottom) Change in thermal
uplift as a function of time. The hatched area is the approximate range in dates for the Marquesan volcanoes
[Duncan and McDougall, 1974; Caroff et al., 1999].
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MCNUTT AND BONNEVILLE: MARQUESAS SWELL
background flux, and the possibility of heat
transport via hydrothermal convection [Harris
et al., 2000] could mean that the total heat flux
is much higher than that. Although it is difficult
to completely rule out the model with the entire
underplating being a melt body, the apparent
lack of appreciable (>100 m) subsidence and of
a large heat flow anomaly (>2 Qo) tends to
support the more modest proposal that the
underplating unit is interlayered melt and mantle material.
[27] This model can be compared to that of
Phipps Morgan et al. [1995], who also proposed
a chemical origin for hotspot swells. Their model, applied to the Hawaiian swell, produces 1200
m of swell relief, with roughly equal contributions from crustal underplating, the transient
thermal effect of the crustal underplating, and a
much deeper low-density depleted mantle residue from melting. This last 90-km-deep component appears to be required for Hawaii in order
to explain the 60-km compensation depth for the
swell [McNutt and Shure, 1986]. If the entire
compensation were chemical and thermal effects
immediately beneath the Moho, the geoid signature would be much smaller. For the Marquesas, the melt fraction extracted from the mantle
may be simply too small to create a substantial
volume of low-density depleted mantle at the
base of the lithosphere.
[28] The simplest explanation remains that the
Marquesas swell is supported within the uppermost mantle just beneath the Moho by crustal
underplating. The fact that the underplating
produces a swell much broader than the footprint of the surface volcanoes may arise from
gravitational spreading of slowly cooling ponds
of magma trapped at the Moho density discontinuity. If this model is correct, its implications for the thermal history of the Marquesas
Islands are quite different than if the swell
support is a deep thermal anomaly at the base
of the lithosphere. For example, it would pre-
1999GC000028
dict that there should only be a short-term,
transient heat flow anomaly over the Marquesas swell. It would also predict a small-amplitude, transient subsidence of the island chain
followed by much slower subsidence at a rate
characteristic of 50±60 Ma seafloor. The subsidence curve would have a vertical offset
caused by the buoyant material just beneath
the crust.
[29] The question remains as to whether the
arguments for support of the Marquesas swell
by a chemically differentiated body lying below the Moho apply to any other hotspot
swells. An underplated body with similar dimension and seismic velocity to that of the
Marquesas has been imaged beneath the Hawaiian swell [Watts and ten Brink, 1989], and
Phipps Morgan et al. [1995] predict that about
two thirds of the swell relief is caused by its
chemical and thermal effects. Other likely candidates for Marquesan-like swell support are
the Canary Islands, with little geoid evidence
for a deeply compensated swell [Filmer and
McNutt, 1989], and the Bermuda Rise, which
has had negligible subsidence over the past
30 ± 40 Ma [Aumento and Ade-Hall, 1973;
Reynolds and Aumento, 1974]. Although we
would not promote the underplating model as
an explanation for all swells, it may be an
important contribution to the geophysical
anomalies associated with hotspots. If so, estimates of the buoyancy flux from hotspots
based on the size of the swell [Davies, 1988;
Sleep, 1990] would require downward revision.
The origin of the geoid lineations (Figure
6) trending perpendicular to the fracture zone
after correcting for the Marquesas Islands and
swell is something of a puzzle. Given that they
trend in the direction of fossil spreading, it is
likely that this geoid signature represents
anomalies frozen into the lithosphere at the
time that the crust was created, as opposed to
present-day dynamic processes that would
[30]
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MCNUTT AND BONNEVILLE: MARQUESAS SWELL
more likely trend in the direction of absolute
plate motion. Their wavelength, 650 km, corresponds to the depth extent of the upper
mantle. Thus it is tempting to imagine that
the geoid stripes represent variations in the
magma supply to this spreading segment of
the Pacific-Farallon ridge caused by convective
instabilities in the upper mantle. However, this
scenario is unlikely for a number of reasons.
First of all, given that the ridge is moving with
respect to the mantle, there is no guarantee that
the anomalies so produced would have the
same trend as the ridge (e.g., the locus of
excess magma could migrate along the ridge
segment with time) or the same wavelength as
the upper mantle. Second, if the variations in
magma supply produced isostatically compensated variations in crustal thickness, either the
corresponding depth anomalies would be very
large, or the compensation depth would have to
be much deeper than the oceanic Moho. Detailed plots of the bathymetry after swell removal show no structures correlated with the
geoid undulations (positive or negative) down
to the 100-m level. There is a slight hint that
the geoid high along the western edge of the
grid is correlated with a 100- to 200-m bathymetric low, but since the lithosphere is older
there, the deeper seafloor may be completely
unrelated to the 650-km geoid undulations. The
uplifted portion of the Marquesas Fracture
Zone Ridge between 1368 and 1388W lies at
the intersection of the eastern geoid high, but
the western high intersects the fracture zone at
one of its deepest points. Thus whatever the
source for the geoid signal, it does not appear
to produce a noticeable or consistent bathymetric effect.
[31] It is intriguing to note, however, that the
Marquesas hotspot began producing surface
volcanoes as it passed through the geoid low.
The trend of the chain departs significantly
from that predicted by absolute motion of the
Pacific plate, as though the eruptions were
1999GC000028
trying to remain within the geoid trough despite
the fact that the hotspot source was located
several hundred kilometers farther to the east.
As the position of the hotspot moved to lie
within the geoid high, the Marquesas hotspot
ceased forming large volcanoes. A very speculative explanation for this behavior could be
that the geoid lineations represent topography
on the base of the lithosphere, with the geoid
lows corresponding to thinner lithosphere that
is more easily penetrated by, and/or forms a
structural trap for hotspot magma. However,
this hypothesis fails to explain how topography
on the base of the lithosphere would fail to
have a surface expression, or the suggestive
correlation between the geoid lineations and
heat flow (Figure 6). While there is no obvious
relationship between heat flow and age of the
lithosphere or distance from the center of the
Marquesas Islands, the two very low heat flow
values do fall within the residual geoid low, and
the normal and high values were collected on
the highs. If so, the geoid low is unlikely
thinned lithosphere, and the position of the
Marquesas Islands within a cold geoid low is
even more puzzling.
5. Conclusions
[32] We propose that the Marquesas swell is
isostatically compensated by a chemically differentiated magmatic body that ponded just
below the Moho. When we account for the geoid
effects of the surface volcanoes, their flexural
compensation, the underplating load at the base
of the crust, and its surface expression, there is no
remaining geoid signal with the orientation of
either the Marquesas hotspot chain or the direction of absolute Pacific plate motion. The parameters in our model, such as the elastic thickness
of the plate, the densities of the units, the location
and height of the volcanoes, the depth of the
moat, and the extent of the underplating body,
were already well determined from previous
gravity, multibeam bathymetry, seismic reflec-
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MCNUTT AND BONNEVILLE: MARQUESAS SWELL
tion, and seismic refraction surveys. Therefore
we had no free parameters to ``adjust.''
[33] For those who would still argue that there is
deeper thermal swell effect, the challenge is
now to find evidence of it in either the geoid
or the bathymetry, given that the known shallower effects remove all signal with a hotspot
orientation. For the Marquesas, and perhaps
Hawaii as well [Phipps Morgan et al., 1995],
the direct thermal and mechanical effect of an
upwelling plume is not evident. To the extent
that underplating exists at other volcanic chains,
the timing and amplitude of subsidence and heat
flow anomalies can be quite different as compared with models that produce swells via
thermal anomalies at the base of the lithosphere.
[34] Fracture-zone-normal lineations in the Marquesas geoid remain after correcting for the
volcanoes, their compensation, the crustal underplating, and its surface uplift. The origin of
these 1-m-high, 650-km-wavelength lineations
is still unexplained but appears to have influenced the position of the Marquesas Islands.
Acknowledgments
[35] This work was funded by NSF OCE-9996270, the
David and Lucille Packard Foundation, and the French
Ministry of Education, Research, and Technology. The
original idea to calculate the topographic effect of the
crustal underplating arose from conversations with M.-A.
Gutscher. We thank Dave Clague, Geoff Davies, Mike
Gurnis, Garrett Ito, Jim Natland, and Paul Wessel for
comments on the manuscript and insight into its implications.
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