1
The Evolution of Large Shield Volcanoes on Venus
Robert R. Herrick
Lunar and Planetary Institute, Houston, Texas
now at Geophysical Institute, University of Alaska Fairbanks
Josef Dufek
University of Chicago, Chicago, Illinois
now at University of Washington, Seattle, Washington
Patrick J. McGovern
Lunar and Planetary Institute
Submitted to: Journal of Geophysical Research – Planets
January 23, 2006
Correspondence Author:
Robert R. Herrick
Geophysical Institute
University of Alaska Fairbanks
903 Koyukuk Dr.
Fairbanks, AK 99775-7320
2
Abstract. We studied the geologic history, topographic expression, and gravity signature of
twenty-nine large Venusian shield volcanoes with similar morphologies in Magellan SAR
imagery. While they appear similar in imagery, sixteen have a domical topographic expression
and thirteen have a central depression. Typical dimensions for the central depression are 150 km
wide and 500 m deep. The central depressions are probably not calderas resulting from collapse
of a shallow magma chamber, but instead are the result of a corona-like sagging of a previously
domical volcano. The depressions all have some later volcanic filling. All but one of the
central-depression volcanoes have been postdated by geologic features unrelated to the volcano,
while most of the domical volcanoes are at the top of the stratigraphic column. Analysis of the
gravity signatures in the spatial and spectral domains shows a strong correlation between the
absence of postdating features and the presence of dynamic support by an underlying plume. We
infer that the formation of the central depressions occurred as a result of cessation of dynamic
support. However, there are some domical volcanoes whose geologic histories and gravity
signatures also indicate that they are extinct, so sagging of the central region apparently does not
always occur when dynamic support is removed. We suggest that the thickness of the elastic
lithosphere may be a factor in determining whether a central depression forms when dynamic
support is removed, but the gravity data is of insufficient resolution to test this hypothesis with
admittance methods.
3
1. Introduction
There are numerous large shield volcanoes on Venus whose flows cover an area more than
500 km in diameter. Venusian volcanoes of this size are generally accepted as forming from hot
spots that are or were located over mantle upwellings [e.g., Head et al., 1992; Solomon et al.,
1992; Phillips and Hansen, 1994; Smrekar et al., 1997]. In contrast, a variety of explanations
have been put forth for the origin of the volcano-tectonic structures known as coronae. Coronae
are quasicircular structures that typically have a raised rim superimposed by an annulus of
concentric fractures or ridges. The largest coronae are comparable in size to the large shield
volcanoes. Explanations for coronae origins include: coronae result from runaway partial
melting in the uppermost mantle initiated by modest spreading [Tackley and Stevenson, 1991,
1993]; coronae are caused by plumes originating from a mid-mantle layer [Stofan et al., 1992];
coronae are caused by breakup of a mantle plume head [Stofan et al., 1995]; coronae form from
plumes interacting with thin lithosphere and volcanoes from plumes interacting with thick
lithosphere [McGill, 1994, 1998; McGovern and Solomon, 1998]; coronae are caused by
detached diapirs, perhaps followed by retrograde subduction and or delamination [Janes et al.,
1992; Sandwell and Schubert, 1992; Koch and Manga, 1996; Smrekar and Stofan, 1997; Jellinek
et al., 2002]; and coronae are formed by small, long-lived plumes from mid-mantle depths that
evolve to delaminate the lithosphere, while shield volcanoes result from large, hot plumes from
the core-mantle boundary [Smrekar and Stofan, 1999].
More recent work [Stofan et al., 2001] has refined the corona database and expanded it to
include 'Type 2' coronae, features with minimal surface fracturing but a topographic signature
typical of coronae. A series of papers also analyzed the distribution of coronae in the expanded
4
database and their gravity signatures [Glaze et al., 2002; Smrekar et al., 2003; Smrekar and
Stofan, 2003]. Those papers attribute the variety of corona morphologies and topographic
shapes primarily to variations in crustal and lithospheric density and the complex interaction
between varying lithospheric properties and a transient plume. The presence or absence of a
depleted mantle layer and possible delamination have significant effects on coronae morphology
and evolution. Initial plume diameters for coronae are interpreted as confined to a narrow range,
and elastic lithospheric thickness is not interpreted to be an important factor in determining the
presence or evolution of coronae.
Other recent work [Herrick, 1999; Johnson and Richards, 2003] has taken a more holistic
approach towards explaining the presence of coronae in relation to a global convection scenario.
The basic idea first advanced in Herrick [1999] and elaborated on in Johnson and Richards
[2003] is that in the Venusian convective regime transient small-scale upwellings coexist with
larger-scale, longer-lived upwellings and downwellings. The small-scale upwellings avoid
major downwelling areas and are 'swept up' if they occur near major upwellings. In this view
mantle upwellings range in form from small, transient, diapiric features to large, long-lived, 'firehose' features. Typical coronae result from the former and regions with one or more large shield
volcanoes result from the latter. Recent physical modeling experiments have indicated that an
absence of plate tectonics, or the prevalence of stagnant lid convection, reduces the typical
viscosity contrast across the base of the mantle thermal boundary layer and makes transient
corona-forming thermals more common on Venus than on Earth [Jellinek et al., 2002]. Other
possibilities arise if both compositional and thermal plumes coexist on a planet. Given the lack
of geochemical data available to test the idea of compositional plumes on Venus, we do not
consider them within this paper.
5
A key issue in determining the origin of coronae is whether or not coronae and shield
volcanoes of similar size are related features. Do the two classes of features result from
variations in a single process, for example plume strength or lithospheric thickness; or do
coronae and volcanoes have separate origins, such as plumes from different boundary layers?
There are a number of features that are listed in compiled databases for both coronae [Stofan et
al., 2001] and volcanoes [Crumpler et al., 1997]. Most of these dually-listed features have
extensive volcanic flow aprons emanating from a central location (hence the volcano
designation) but are topped with a wide, flat plateau that is surrounded by a raised and/or faulted
annulus (hence the corona designation). This dual classification could result from the grouping
together of features with different origins by the database-compiling groups. Alternatively, these
dually classified features could represent transitional forms between shield volcanoes and
coronae [e.g., McGill, 1994, 1998; McGovern and Solomon, 1998].
Research we have conducted over the past few years has led to development of the following
hypothesis: During the extinction phase of volcano formation, the central region of a large
shield volcano sags when dynamic support from a long-lived mantle plume is removed; the
width of the sagged region reflects the width of the former plume stem. This hypothesis was
developed in part based on results of a regional geological-geophysical study of the KunhildEreshkigal region on Venus [Herrick and McGovern, 2000]. Kunhild and Ereshkigal are similar
in morphology to Sif Mons and Sappho Patera, respectively, two structures previously
interpreted to be located over broad mantle upwellings and sites of geologically recent (or
active) volcanism [Grimm and Phillips, 1992; Senske et al., 1992; McGill, 1994]. However,
unlike Sif and Sappho, a low gravity signature and photogeologic mapping indicate that the
region is currently extinct and not located over a mantle upwelling [Herrick and McGovern,
6
2000]. The terminal stage of both features involves a sagging of the central region. For
Ereshkigal the process is recognizable as an example of the final stage of corona formation, and
the central collapse is attributed to loss of dynamic support resulting from the spreading and
dissipation of a hot thermal diapir [e.g., Squyres et al., 1992; Koch, 1994; Musser and Squyres,
1997]. The abundant volcanism associated with Ereshkigal has led to Ereshkigal's classification
as both a volcano and a corona. McGill [1998] made a similar interpretation that Sappho was a
shield volcano that evolved into a corona, and surmised that it is in its waning stages of activity.
Kunhild's appearance and overall topographic shape indicate it is a typical large Venusian shield
volcano, yet its terminal stage involves a central sagging that appears very similar in nature to
what occurred at Ereshkigal.
Because of its unusual size and morphology, it is unlikely that Kunhild's central depression is
a caldera in the traditional sense of the word. Calderas approaching 100 km in diameter on Earth
are associated with explosive ignimbrite eruptions; examples of basaltic volcanoes with
comparable-sized calderas are the largest Martian shield volcanoes. For a typical volcanic
caldera on the Earth or Mars, the collapse of the surface over a shallow magma chamber
produces circumferential steep-sided faults with significant throw [e.g., Lipman, 2000]. In
contrast, Kunhild's depression is surrounded by a gradual upslope to a radially-fractured annulus,
and in places it appears that preexisting flows now drape over this annulus. This suggests that the
cause of Kunhild's central depression was deep and that the surface response was a distributed
downwarping of the summit region. A similar appearing radially fractured rim embayed by
interior volcanic deposits has been observed in coronae of similar size [e.g., Eastern Eistla Regio
in Smrekar and Stofan, 1999]. While sag features, rather than discrete faults, have been
observed at some terrestrial volcanoes, in those cases the diameter of the sagged region is at
7
most a few tens of kilometers and still consistent with shallow withdrawal of magma [e.g.,
Walker, 1984].
Here we test the hypothesis of Herrick and McGovern [2000] with a survey of large shield
volcanoes on Venus. Our approach is to analyze the imagery, topography, and gravity signatures
of a number of similar-appearing Venusian shield volcanoes. In its basic form, our hypothesis
predicts that extinct shield volcanoes should have a central sagged region while active volcanoes
should not. Thus, shield volcanoes with a broad central low would be expected to correlate with
the presence of unrelated postdating geologic features and with a gravity signature consistent
with isostatic compensation. Active volcanoes are expected to be at the top of the local
stratigraphic column and have a gravity signature consistent with some level of dynamic support
from the mantle.
2. General Characterization
On the basis of their appearance in Magellan imagery, we selected 29 shield volcanoes for
detailed analysis. Our goal was to select a population that are generally similar in appearance.
The volcanoes were subselected from the volcano database of Crumpler et al. [1997] because
they appear to be broadly domical, are axisymmetric, and have radial flows exceeding 500 km in
diameter. Our data set includes a few volcanoes listed as between 300 and 500 km in diameter in
the Crumpler et al. [1997] for which we determined that the flow apron diameter had been
underestimated. Our survey represents about half of the 500+ km diameter volcanoes in the
Crumpler et al. [1997] database. The selected volcanoes do not appear to have any steep slopes
or sharp peaks, and they are not cut by major rifts (although they may superpose a rift). Table 1
identifies the 29 volcanoes along with their basic characteristics. Location and flow apron
8
diameter are taken from Crumpler et al. [1997]. Height is listed relative to the typical elevation
of the surrounding terrain. Topographic width is characterized by two measures: the width of
the volcano at half of its maximum elevation, and the width of the volcano at an elevation 500 m
above the surrounding terrain. All topographic measurements were made from perpendicular
profiles across each volcano using the gridded Magellan altimetry data (GTDR).
2.2 Topography
Using Magellan radar altimetry (~10 km resolution) the volcanoes were characterized by the
presence or absence of a central depression of width greater than 100 km (Table 1). Sixteen of
the features had no significant topographic depression and a broad domical shape (labeled "d"
for domical in Table 1, column 4). Ten of the volcanoes had a broad central depression (labeled
"c" for central depression). Three of the features appear to have broad central depression with
volcanic edifices of significant size in their interiors (labeled "ce" for central depression with a
large interior edifice). Figure 1 shows some examples of the different categories of volcanoes
with topographic profiles.
2.3 Postdating Features
There is no available information that can be used to estimate an absolute age for a Venusian
volcano. Furthermore, there are too few craters on Venus to use crater counting techniques to
evaluate the relative ages of the volcanoes. However, a crude comparison of the relative ages of
the different categories of volcanoes can be made by observing what features, if any, postdate
each of the volcanoes. If members of one group of volcanoes consistently have a number of
geologic features that postdate the volcano, then it is reasonable to interpret this group as
containing many extinct volcanoes. In contrast, if members of the group are all at the top of the
9
stratigraphic column, then the possibility exists that many of these volcanoes are currently or
recently active.
We have used basic photogeologic mapping techniques to determine if there are any geologic
features that postdate the volcano that are unrelated to its formation. Our photogeologic analysis
was performed using the full resolution Magellan imagery (resolution ~100 m). The nature of
the features that postdate the volcano are as follows (Table 1, column 8): impact craters (C),
embayment by flows from another volcano (E), wrinkle ridges (R), fractures (F), and lineaments
(L) of unknown origin. Figure 2 shows examples of these relationships. Impact craters with
radar-dark floors are likely to be partially flooded by volcanic flows [Herrick and Sharpton,
2000], so only impact craters with radar-bright floors and ejecta blankets that appeared intact
were considered as a feature post-dating the volcano.
2.4 Isostatic Anomaly
Using the spherical harmonic gravity field of Konopoliv et al. [1999] and the topography of
Rappaport et al. [1999] we have calculated the global isostatic anomaly. The resolution of the
Magellan-derived gravity varies across the planet, but most of the globe can be considered
reliable to degree and order 75 based on the degree strength map of Konopoliv and Sjogren
[1996]. The isostatic anomaly was calculated globally (to degree and order 75) by subtracting
the gravity signal attributable to the Airy-compensated topography from the free-air gravity
anomaly. We assumed a crustal thickness of 30 km and a crustal density of 2900 kg m-3 [Grimm
and Hess, 1997]. Table 1 gives the value of the anomaly at the center of the structure and
whether it appears correlated with the location of the volcano, and Figure 3 shows sample
profiles of the isostatic anomaly for Sif (correlated) and Kokyanwuti (uncorrelated). A
10
significant, correlated positive anomaly (tens of mGals) can be interpreted as indicative of partial
support of the volcano by heated mantle material. Significant flexural support of the volcano is a
possible alternative explanation. A strong correlation between absence of postdating features
and a significant isostatic anomaly favors the interpretation that the isostatic anomaly results
from dynamic support. We interpret an anomaly with absolute magnitude near zero as indicative
of crustal support by Airy isostasy (Pratt isostasy is also an acceptable interpretation). The
results are insensitive to the exact choices of crustal density and thickness.
2.5 Central depressed regions
A typical size for the central depressed region is on the order of 150 km across and a few
hundred meters deep (Table 2). All of the depressed central regions appear to have a geologic
history similar to that described for Kunhild in Herrick and McGovern [2000]. The central
depressed region in the studied volcanoes appears to have formed through sagging of the
topography. Circumferential normal faulting occurs in only a few of the volcanoes, and the
throw is minimal. Several volcanoes have finely spaced radial fractures along the topographic
rim of the depression, and in some cases volcanic flows can be seen that drape over the rim.
Partial volcanic filling has occurred in all of the volcanoes with a depressed central region. A
spectrum of volcanic edifice types and sizes appear in the interiors of the depressed regions,
ranging from a few small shields to features that are large enough to fill almost the entire
depressed region. There are end member volcanoes that have topographic breaks in slope and
morphologic remnants of a depressed central region, but no significant central depression
currently exists. An interpretation we favor is that a central sagged region existed that has been
nearly filled by post-sagging edifices, and these volcanoes are labeled "ce" in Table 1.
11
2.6 Discussion
The statistical measurements of elevation and width of the volcanoes combined with our
general observations confirm that we have selected a family of reasonably similar volcanoes.
There are some rather obvious correlations between the different columns of Table 1. Table 3
shows the mean and median isostatic anomaly for the different volcano types. Volcanoes with
no post-dating features are clearly most likely to have a large isostatic anomaly indicative of
dynamic support, and volcanoes with abundant post-dating features generally have gravity
signals consistent with isostatic support of the edifice. This correlation is expected and is
consistent with the idea that a large isostatic anomaly indicates dynamic support from a mantle
plume beneath a currently or recently active volcano.
All of the volcanoes with central depressions are consistent with isostatic support, and all but
one have clear post-dating features. Most of the domical volcanoes have large isostatic
anomalies and no post-dating features, but several have post-dating features and are consistent
with isostatic support. What is not easily quantifiable for placement in a data table is the
pervasiveness of postdating features. For domical volcanoes a postdating feature is likely to be a
few fractures or folds deforming some of the distal flows, while the volcanoes with central
depressions are often permeated by tectonic deformation. Nevertheless, it can be argued that at
least some of the domical volcanoes are extinct features.
3. Two-layer dynamic inversion
3.1 Technique
To place the gravity signature associated with each volcano in a form that allows more
12
quantitative interpretation, we employed both spatial and spectral domain approaches. In this
section we discuss results from a spatial-domain inversion performed using the technique
described in Herrick and Phillips [1992]. The approach is to use two constraints, the gravity and
topography, to solve uniquely for surface densities on two subsurface layers. Stated a different
way, it is always possible to match the free-air gravity signal by compensating the topography
with sheets of mass anomalies placed at two depths. In the particular model we use here, we
assume that the topography is supported by Airy isostasy at the base of the crust and by upper
mantle convective flow. If a nominal crustal thickness is assigned, then Airy isostasy at the base
of the crust can be represented by a sheet of varying surface density placed at a depth equal to
the assigned crustal thickness. It is reasonable to assume that the effect of mantle convection on
the gravity and topography is dominated by the flow, and corresponding mass anomalies, in the
upper mantle. We can represent these anomalies by a sheet of varying surface density placed at
a depth somewhat below the base of the conductive lid (i.e., the thermal lithosphere) for the
convective system. If a viscosity structure for the mantle is specified and we assume that mantle
convection is being driven by that mass sheet, then we can solve for the surface effect on the
gravity and topography using the technique described in Richards and Hager [1984; for further
description see also Hager and O'Connell 1979, 1981; Hager and Clayton, 1989].
We make all of the same model assumptions used in Herrick and Phillips [1992]. Here we
restate the most important features of the model. The crustal density is assumed to be 2900 kg
m-3, and the mantle density just below the crust is assumed to be 3300 kg m-3. We have assumed
that the mantle is an incompressible Newtonian viscous fluid with a no-slip condition at the
surface and a free-slip condition at the core-mantle boundary. We assume an isoviscous mantle.
The base of the crust is assumed to be at 30 km depth, and the perturbing shell in the mantle is
13
placed at 200 km depth. The model parameters are the favored values found in Herrick and
Phillips [1992]. That work experimented with a variety of shell depths and viscosity models.
Our assumed nominal crustal thickness of 30 km is also consistent with typical estimates of the
Venusian crustal thickness [e.g., Grimm and Hess, 1997]. A depth of 200 km for the lower,
convective shell is consistent with the most important mantle perturbations being located in the
upper mantle below a conductive lid of 100 km thickness, a typical estimated thermal lithosphere
thickness [e.g, Phillips and Malin, 1983; Phillips et al., 1997; Simons et al., 1997; Brown and
Grimm, 1999]. We carried out the inversion through harmonic degree and order 90, with a cos2
taper of the coefficients from degrees 60 through 90. Because we are comparing features in the
spatial domain, it is important to perform the inversions using the same number of harmonics.
The degree strength of the gravity field is greater than 60 for all but one of the features (Table 1),
so we feel we are not introducing significant noise into the inversions.
The choices of parameters for the inversion are certainly debatable and largely unconstrained.
However, they are adequate for the purposes of broad comparisons and general hypothesis
testing.
3.2 Results
Figure 4 shows typical results of the inversions. Table 4 shows the surface density for each
shell over the center of each volcano and identifies whether the volcano is distinguishable from
its surroundings on the given surface. Figure 5 graphically summarizes the results. The domical
volcanoes with no post-dating features all had modest to significant negative surface densities on
the lower, mantle convection shell that are consistent with dynamic support by upward moving
buoyant material in the mantle. A number of additional assumptions must be made to translate
14
surface densities into meaningful quantities. For example, if we wish to interpret the densities in
terms of a thermal plume, we must make assumptions about the thickness of the perturbing layer
and the coefficient of thermal expansion. If we assume a 200 km thick perturbing layer and a
nominal upper mantle density of 3300 kg m-3, then a surface density anomaly of -106 kg m-2
translates into a density change in that layer of 5 kg m-3 , or about 0.15%. If we assume a
coefficient of thermal expansion of 3 x 10-5 K-1, then this density anomaly represents a 50 K
temperature anomaly.
Most of the domical volcanoes with no post-dating features have surface density anomalies on
the mantle convection shell that have maximum negative values of -2 x 106 to -6 x 106 kg m-2.
Using the above assumptions, this implies density changes of 0.3% to 0.9%, or central plume
temperatures 100 – 300 K above the ambient mantle. These values are similar to estimates from
seismic tomographic studies for Earth at comparable resolutions over locations of suspected
mantle plumes [see review by Nataf, 2000], and they are consistent with estimates of excess
plume temperatures derived from geochemical data [e.g., Schilling, 1991; also see review by
Sleep, 1992]. Using these assumptions, Figure 6 shows an example of the surface densities on
the lower subsurface shell converted to temperature for Ushas Mons.
Most of the volcanoes have similar modest negative anomalies,-1 x 106 to -4 x 106 kg m-2, on
the upper subsurface shell (representing perturbations of the crust-mantle boundary) that are
consistent with a crustal root beneath the volcano [Table 4]. For example, if the crustal density
is 2900 kg m-3, then isostatic compensation of a 1 km high volcano would require a surface
density of -2.9 x 106 kg m-2 on the upper subsurface shell. Some of the domical volcanoes with
no post-dating features have positive anomalies on the upper subsurface shell. This may
represent some amount of crustal thinning over a putative plume. Alternatively, it could
15
represent some level of flexural support.
Consistent with the idea that the volcanoes with postdating features are geologically inactive
and no longer located over an upwelling plume, most of these volcanoes have surface densities
on the lower subsurface shell that are small in magnitude. Some of these features, however, do
have significant positive anomalies on the lower subsurface shell. This can occur if the nominal
crustal thickness is assumed to be deeper than its true value. For these volcanoes, the positive
anomaly on the lower subsurface shell is eliminated or substantially reduced if the upper shell is
placed at 15 km depth instead of 30 km [Figure 7]. Alternatively, there could be a real positive
density anomaly in the upper mantle beneath some of the volcanoes. Possibilities include a
denser than normal lower crust or a thermal downwelling. The positive anomaly is also
essentially eliminated if the higher order harmonics are not used [Figure 8]. Among the
volcanoes with postdating features, it appears that those with a depressed central region are more
likely to have a positive density anomaly on the lower subsurface shell that correlates with the
location of the volcano, but there are enough exceptions that we cannot confidently assert that
this is a characteristic of those volcanoes. For example, from the data one could argue that a
positive lower-surface anomaly that correlates with the volcano always occurs, but it is only
observable in areas where the gravity field has a degree-strength greater than ~85; i.e., the
positive anomaly is only observed if there is no noise in the data for harmonic degrees 60 to 90.
4. Spatio-spectral localization
4.1 Technique
Utilizing the method developed and presented in Simons et al. [1997], we determined the
RMS amplitude of the gravity, topography, and admittance (ratio of gravity to topography)
16
spectra for areas centered on each of the 29 volcanoes. The technique is to multiply the spherical
harmonic representation of the planet by a spherical cap centered on the location of each
volcano, and then analyze the resulting spectra. This is the spherical harmonic version of
multiplying a one-dimensional function by a boxcar (or some other finite, symmetric function)
and then evaluating the power spectra. The technique is a convolution with a broad filter in the
spectral domain, and the resulting spectra can be considered valid only where the filter is
convolved with valid data. Consequently, the localized spectra contains usable data only over
the spectral range of the initial data minus the width of the filter. The localized spectra can also
be thought of as smoothed over the width of the filter. There is a trade-off between the width of
the cap in the spatial domain and the width of the filter in the spectral domain. A narrow cap,
which better localizes a feature, generates a smoothed spectra with usable data over a small
range of harmonics. Because the ultimate goal is to interpret the spectra in terms of a surface
loading model, one can restate the problem as smaller caps providing less data to use in creating
a robust spectra.
The Simons et al. [1997] paper suggested two simple choices for caps. The choice they used
was a cap whose width was a set number of "wavelengths" for each harmonic degree. In the
spatial domain the cap varies in width for each harmonic degree, meaning the spectra is more
localized for high-degree harmonics versus low-degree harmonics. In the spectral domain this
results in convolution with a filter that broadens for higher harmonic degrees (once again, a
trade-off between spatial and spectral resolution). Another suggested choice was a spherical cap
of a fixed horizontal width equal to one wavelength of zonal harmonic degree Lwin (width ≅ 2 π
Rplanet / Lwin ). This results in convolution with a fixed width filter in the spectral domain. A
familiar analog is that multiplying a 1-dimensional function by a boxcar represents convolution
17
with a sinc function in the spectral domain. The first choice, the varying-width cap, is wellsuited for evaluating global patterns of variations in admittance (the focus of the Simons et al.
[1997] paper), while the fixed-width cap is conceptually easier to visualize when comparing the
spectra of individual features. The fixed-width cap was successfully employed by McGovern et
al. [2002] to evaluate the admittance spectras of selected features on Mars, and it was also used
in the investigation of the Kunhild-Ereshkigal region on Venus [Herrick and McGovern, 2000].
In this work we use a fixed-width cap, and the usable range of harmonic l in the windowed fields
is Lwin ≤ l ≤ (Lobs – Lwin), where Lobs is the maximum degree of observation. For admittances on
Venus the gravity data is the limiting factor, and we use the degree strength of the gravity field
[Konopoliv and Sjogren, 1996] as an estimate of Lobs.
It is also possible to evaluate the spectra in terms of specific models of support of the
topography by Airy isostasy and elastic flexure. We utilize the model for loading of a thin
elastic shell described in McGovern et al. [2002], which is based on models presented by
Turcotte et al. [1981] and Kraus [1967]. The general technique is to generate a global free-air
gravity model using the global topography compensated with a specified set of parameters for
the crust and lithosphere. The gravity model is then windowed in an identical manner to the
actual gravity to produce a model curve for admittance in a given region. We chose simple
models with top-loading only, a load density and crustal density of 2900 kg m-3, and an upper
mantle density of 3300 kg m-3. The elastic lithospheric thickness and the nominal crustal
thickness were varied in the models from 0 to 40 km.
4.2 Results and Discussion
In our initial spectral localization we used a value of Lwin = 15 for a cap width of 2500 km.
18
We feel this is a reasonable compromise between isolating the large volcano and measuring a
broad spectral range. Figure 9 shows several examples of the admittance spectras, and
admittance and correlation values at several different harmonic degrees are shown in Table 5.
The plots in Figure 9 also show model admittances for the topography compensated with crustal
thicknesses of 20 and 40 km, and elastic lithospheric thicknesses of 10 and 30 km. There are a
few general patterns in the data that we feel confident in identifying. Admittances for the
domical volcanoes without any postdating features are significantly higher than the admittances
for the volcanoes with depressed central regions up thru degree 40. Also, for degrees 30 through
50 the correlation between gravity and topography is significantly lower for volcanoes with
depressed central regions than domical volcanoes with no post-dating features. A decorrelation
between the gravity and topography coefficients can be thought of in the spatial sense as
different features within the cap requiring different depths of compensation. Alternatively, one
can envision density anomalies deep enough that their effect on surface topography is minimal.
It is more difficult to assess whether there are significant differences between the different
classes of volcanoes that have postdating features. We calculated means for a variety of
quantities after dividing the volcanoes into three groups: domical volcanoes without postdating
features, domical volcanoes with postdating features, and all other volcanoes. These summary
plots are shown in Figure 10. Because the volcanoes of Bell Regio are in close proximity to
each other but in different groups, they are not included in the summary plots. Hathor and Innini
are in close enough proximity that their spectras were almost identical, and only Hathor's spectra
is shown in the summary plots. In neither the summary plots nor the individual plots can we see
a substantive difference between the spectra of the different types of volcanoes with postdating
features.
19
Our general interpretation of the spectral data can be summarized as follows. As a family, the
domical volcanoes without postdating features are partially dynamically supported by a mantle
plume whose location correlates with the location of the volcano. The result is high admittances,
particularly at longer wavelengths (up to degree 40), and a strong correlation between gravity
and topography. The remaining volcanoes have admittances at long wavelengths that are
modestly higher than can be accounted for by lithospheric compensation, but the nonlithospheric
component contributes minimally to the regional gravity and topography signature. Without
regional gravity and topography being dominated by a single plume, gravity-topography
correlation drops off significantly above about degree 30 or 40. No clear distinction can be
made in the spectral domain between different types of volcanoes with postdating features. The
spectral data are not interpretable for the purpose of discerning the nominal thickness of the crust
or elastic lithosphere in the regions encompassing the volcanoes.
5. Conclusions
There are some conclusions that we feel can be made confidently, and there are others that are
more tenuous. Our results strongly suggest that shield volcanoes with central depressed regions
are not currently active features and are not receiving dynamic support from a mantle plume.
These volcanoes are consistently post-dated by unrelated geologic features, and there is no
indication in the gravity data that they are underlain by a mantle upwelling. We feel confident
that many of the domical volcanoes are recently active and are located over a mantle plume.
They are at the top of the stratigraphic column, and they have gravity signatures consistent with
dynamic support in both the spatial and spectral domains. The morphologies of the depressed
central regions indicate that the central depression forms from sagging of the center of a domical
20
volcano, and modest volcanism then later partially fills the depression.
Based on these observations, it is tempting to conclude that the evolutionary sequence of large
shield volcanoes on Venus always involves a terminal stage of corona-like sagging of the interior
when the mantle plume feeding the volcano dies out. This interpretation would suggest that
plume duration and the associated amount of constructional volcanism is an important factor in
determining whether a corona or a shield volcano forms at a given location. However, there are
some volcanoes that are domical that are best interpreted as extinct features. They are postdated
by unrelated geologic features and have gravity signatures consistent with static support within
the lithosphere. If shield volcanoes always evolve to having a central depression, then the
extinct domical volcanoes must represent an intermediate stage. A possible evolutionary
sequence that parallels an evolutionary sequence proposed for coronae [Smrekar and Stofan,
1997, 1999] would be that the central depression actually results from delamination, and extinct
domical volcanoes represent a phase between upwelling and later delamination. This, however,
is a rather ad hoc explanation and testing its plausibility through numerical modeling is beyond
the scope of this investigation.
We feel the morphologic evidence is compelling that the volcanoes with central depressions
were domical at some point in their history. If the extinct domical volcanoes are at an end stage,
then there must be something different that causes the terminal stage of a shield volcano to vary.
McGovern and Solomon [1998] showed that elastic lithospheric thickness at the time of
formation may be an important factor in determining whether a volcano or corona forms, with
coronae forming over thin lithosphere and shield volcanoes over thick lithosphere. Volcanoes
with central depressions were suggested to represent a transitional state between the two. The
absence of coronae and dominance of shield volcanoes over present-day regions of large mantle
21
upwellings (where the lithosphere is likely to be thinnest) argues against the elastic lithospheric
thickness being the only differentiator between corona versus volcano formation [Herrick, 1999;
Johnson and Richards, 2003]. We suggest here that, if other factors are similar, the thickness of
the elastic lithospheric is important in determining whether a shield volcano has a terminal stage
of central sagging. Unfortunately, we do not feel the gravity data are of adequate resolution for
our sample set to be able to distinguish the elastic lithospheric thickness for the different types of
volcanoes with postdating features.
As we discuss above, we do not favor the interpretation that the central depressions are
calderas. However, if one chooses to view the central depressions as unusually large calderas,
extinction of the volcano still occurs when the underlying plume goes away, but the depression
results from the physical withdrawal of magma rather than the deeper removal of thermal
support. A key difference in the volcano's evolution is that the structure of the central depression
may form early in a volcano's history, but be kept filled until the volcano becomes extinct. The
appeal of this interpretation is its ability to account for the partial to complete interior
embayment of central depressions. In this case, the genetic cause for the central depression in a
large shield volcano is different from that associated with a corona.
In summation, our observations indicate that sagging of the central region is a common
occurrence in the late stages of shield volcano formation on Venus, but this sagging apparently
does not always occur. We suggest that both plume duration and elastic lithospheric thickness
are important factors in determining whether the final product of a mantle upwelling is a corona,
a domical shield volcano, or a shield volcano with a central sagged region.
Acknowledgements. The authors thank Michael Manga and an unidentified reviewer for
useful comments. This is LPI Contribution 1226.
22
23
References
Brown, C. D., and R. E. Grimm, Recent tectonic and lithospheric thermal evolution of Venus,
Icarus, 139, 40-48, 1999.
Crumpler, L. S., and 5 others, Volcanoes and centers of volcanism on Venus, in Venus II, eds. S.
W. Bougher, D. M. Hunten, and R. J. Phillips, pp. 697-756, U. of Arizona Press, Tucson,
1997.
Glaze, L. S., E. R. Stofan, S. E. Smrekar, and S. M. Baloga, Insights into corona formation
through statistical analyses, J. Geophys. Res., 107(E12), 5135, doi:10.1029/2002JE001904,
2002.
Grimm, R. E., and P. C. Hess, The crust of Venus, in Venus II, eds. S. W. Bougher, D. M.
Hunten, and R. J. Phillips, pp. 1205-1244, U. of Arizona Press, Tucson, 1997.
Grimm, R. E., and R. J. Phillips, Anatomy of a Venusian hot spot: Geology, gravity, and mantle
dynamics of Eistla Regio, J. Geophys. Res., 97, 16,035-16,054, 1992.
Hager, B. H., and R. W. Clayton, Constraints on the structure of mantle convection using seismic
observations, flow models, and the geoid, in Mantle Convection: Plate Tectonics and Global
Dynamics, edited by W. R. Peltier, pp. 657-764, Gordon and Breach, New York, 1989.
Hager, and R. J. O'Connell, Kinematic models of large-scale flow in the Earth's mantle, J.
Geophys. Res., 84, 1031-1048, 1979.
Hager, and R. J. O'Connell, A simple global model of plate dynamics and mantle convection, J.
Geophys. Res., 86, 4843-4867, 1981.
Head, J. W. L. S. Crumpler, J. C. Aubele, J. E. Guest, and R. S. Saunders, Venus volcanism:
Classification of volcanic features and structures, associations, and global distribution from
Magellan data, J. Geophys. Res., 97, 13,153-13,199, 1992.
Herrick, R. R., Small mantle upwellings are pervasive on Venus and Earth, Geophys. Res. Lett.,
26, 803-806, 1999.
Herrick, R. R., and P. J. McGovern, Kunhild and Ereshkigal, an extinct hot-spot region on
Venus, Geophys. Res. Lett., 27, 839-842, 2000.
Herrick, R. R., and R. J. Phillips, Geological correlations with the interior density structure of
Venus, J. Geophys. Res., 97, 16,107-16,034, 1992.
Herrick, R. R., and V. L. Sharpton, Implications from stereo-derived topography of Venusian
impact craters, J. Geophys. Res., 105, 20,245-20,262, 2000.
Janes, D. M., and 6 others, Geophysical models for the formation and evolution of coronae on
Venus, J. Geophys. Res., 97, 16,055-16,067, 1992.
Jellinek, A. M., The influence of interior mantle temperature on the structure of plumes: Heads
for Venus, tails for the Earth, Geophys. Res. Lett., 29, 10.1029/2001GL014624, 2002.
Johnson, C. L., and M. A. Richards, A conceptual model for the relationship between coronae
and large-scale mantle dynamics on Venus, J. Geophys. Res., 108(E6), 5058,
24
doi:10.1029/2002JE001962, 2003.
Koch, D. M., A spreading drop model for plumes on Venus, J. Geophys. Res., 99, 2035-2052,
1994.
Koch, D. M., and M. Manga, Neutrally buoyant diapirs: A model for Venus coronae, Geophys.
Res. Lett., 23, 225-228, 1996.
Konopliv, A. S., and W. L. Sjogren, Venus Gravity Handbook, JPL Publication 96-2, Pasadena,
1996.
Konopliv, A. S., W. B. Banerdt, and W. L. Sjogren, Venus gravity: 180th degree and order
model, Icarus, 134, 3-18, 1999.
Kraus, H., Thin Elastic Shells, John Wiley, New York, 1967.
Lipman, P. W., Calderas, in Encyclopedia of Volcanoes, ed. H. Sigurdsson, pp. 643-662,
Academic Press, San Diego, 2000.
McGill, G. E., Hotspot evolution and Venusian tectonic style, J. Geophys. Res., 99, 23,14923,161, 1994.
McGill, G. E., Central Eistla Regio: Origin and relative age of topographic rise, J. Geophys.
Res., 103, 5889-5896, 1998.
McGovern, P. J., and S. C. Solomon, Growth of large volcanoes on Venus: Mechanical models
and implications for structural evolution, J. Geophys. Res., 103, 11,071-11,101, 1998.
McGovern, P. J., S. C. Solomon, D. E. Smith, M. T. Zuber, M. Simons, M. A. Wieczorek, R. J.
Phillips, G. A. Neumann, O. Aharonson, and J. W. Head, Localized gravity/topography
admittance and correlation spectra on Mars: Implications for regional and global evolution,
J. Geophys. Res., 107(E12), 5136, doi:10.1029/2002JE001854, 2002.
Musser, G. S., and S. W. Squyres, A coupled thermal-mechanical model for corona formation on
Venus, J. Geophys. Res., 102, 6581-6595, 1997.
Nataf, H.-C., Seismic imaging of mantle plumes, Annu. Rev. Earth Planet. Sci., 28, 391-417,
2000.
Phillips, R. J., and V. L. Hansen, Tectonic and magmatic evolution of Venus, Ann. Rev. Earth
Planet. Sci., 22, 597-654, 1994.
Phillips, R. J., and M. C. Malin, The interior of Venus and tectonic implications, in Venus, edited
by D. M. Hunten, L. Colin, T. M. Donahue, and V. I. Moroz, pp. 159-214, University of
Arizona Press, Tucson, 1983.
Phillips, R. J., C. L. Johnson, S. L. Mackwell, P. Morgan, D. T. Sandwell, and M. T. Zuber,
Lithsopheric mechanics and dynamics of Venus, in Venus II, eds. S. W. Bougher, D. M.
Hunten, and R. J. Phillips, pp. 1163-1204, U. of Arizona Press, Tucson, 1997
Rappaport, N. J., A. S. Konopliv, and A. B. Kucinskas, An improved 360 degree and order
model of Venus topography, Icarus, 134, 19-31, 1999.
Richards, M. A., and B. H. Hager, Geoid anomalies in a dynamic Earth, J. Geophys. Res., 89,
5987-6002, 1984.
25
Sandwell, D. T., and G. Schubert, Flexural ridges, trenches, and outer rises around coronae on
Venus, J. Geophys. Res., 97, 16,069-16,084, 1992.
Schilling, J., Fluxes and excess temperatures of mantle plumes inferred from their interaction
with migrating mid-ocean ridges, Nature, 352, 397-403, 1991.
Senske, D. A., G. G. Schaber, and E. R. Stofan, Regional topographic rises on Venus: Geology
of Western Eistla Regio and comparison to Beta Regio and Atla Regio, J. Geophys. Res., 97,
13,395-13,420, 1992.
Simons, M., S. C. Solomon, and B. H. Hager, Localization of gravity and topography:
constraints on the tectonics and mantle dynamics of Venus, Geophys. J. Int., 131, 24-44,
1997.
Sleep, N. H., Hotspot volcanism and mantle plumes, Annu. Rev. Earth Planet. Sci., 20, 19-43,
1992.
Smrekar, S. E., and E. R. Stofan, Corona formation and heat loss on Venus by coupled upwelling
and delamination, Science, 277, 1289-1294, 1997.
Smrekar, S. E., and E. R. Stofan, Origin of corona-dominated topographic rises on Venus,
Icarus, 139, 100-115, 1999.
Smrekar, S. E., and E. R. Stofan, Effects of lithospheric properties on the formation of Type 2
coronae on Venus, J. Geophys. Res., 108(E8), 5091, doi:10.1029/2002JE001930, 2003.
Smrekar, S. E., W. S. Kiefer, and E. R. Stofan, Large volcanic rises on Venus, in Venus II, eds.
S. W. Bougher, D. M. Hunten, and R. J. Phillips, pp. 845-878, U. of Arizona Press, Tucson,
1997.
Smrekar, S. E., R. Comstock, and F. S. Anderson, A gravity survey of Type 2 coronae on Venus,
J. Geophys. Res., 108(E8), 5090, doi:10.1029/2002JE001935, 2003.
Solomon, S. C., and 10 others, Venus tectonics: An overview of Magellan observations, J.
Geophys. Res., 97, 13,199-13,256, 1992.
Squyres, S. W., and 6 others, The morphology and evolution of coronae on Venus, J. Geophys.
Res., 97, 13,611-13,634, 1992.
Stofan, E. R., and 6 others, Global distribution and characteristics of coronae and related features
on Venus: Implications for origin and relation to mantle processes, J. Geophys. Res., 97,
13,347-13,378, 1992.
Stofan, E. R., S. E. Smrekar, D. L. Bindschadler, and D. A. Senske, Large topographic rises on
Venus: Implications for mantle upwelling, J. Geophys. Res., 100, 23,317-23,327, 1995.
Stofan, E. R., S. E. Smrekar, S. W. Tapper, J. E. Guest, and P. M. Grindrod, Preliminary analysis
of an expanded corona database for Venus, Geophys. Res. Lett., 28, 4267-4270, 2001.
Tackley, P. J., and D. J. Stevenson, The production of small Venusian coronae by RayleighTaylor instabilities in the uppermost mantle (abstract), Eos, Trans. AGU, 72, 287, 1991.
Tackley, P. J., and D. J. Stevenson, A mechanism for spontaneous self-perpetuating volcanism
on the terrestrial planets, in Flow and Creep in the Solar System: Observations, Modeling
and Theory, eds. D. B. Stone and S. K. Runcorn, pp. 307-321, Kluwer Academic Publishers,
26
Netherlands, 1993.
Turcotte, D. L., R. J. Willemann, W. F. Haxby, and J. Norberry, Role of membrane stresses in
support of planetary topography, J. Geophys. Res., 86, 3951-3959, 1981.
Walker, G. P. L., Downsag calderas, ring faults, caldera sizes, and incremental caldera growth, J.
Geophys. Res., 89, 8407-8416, 1984.
27
Figure Captions
Figure 1. Imagery and topographic profiles for several shield volcanoes (trailing letter
indicates volcano type in table 1): (a) Sif Mons – d, (b) Innini Mons – d, (c) Kunapipi Mons – d,
(d) Chloris Mons – ce, (e) Kokyanwuti Mons - c, and (f) Atanua Mons - c.
Figure 2. Profiles of the isostatic anomaly for (a) Sif Mons – d, and (b) Kokyanwuti - c. See
Figure 1 for locations of profiles.
Figure 3. Examples of the different types of features that postdate Nzambi Corona (top) and
Mielikki Mons (bottom).
Figure 4. Examples of inversions for surface densities on two layers for (a) Ushas Mons - d,
(b) Hathor Mons – d, and Innini Mons - d, (c) Tuulikki Mons - d, and (d) Atanua Mons - c.
Upper layer is isostatic compensation at a depth of 30 km, and lower layer is a mass sheet
driving mantle flow placed at a depth of 200 km. Inversion was carried out through harmonic
degree and order 90, with a cos2 taper from degrees 60 to 90. Contour interval is 106 kg m-2.
Scale bar at image bottom is 500 km in length.
Figure 5. Surface densities on upper and lower subsurface shells (units 106 kg m-2 for):
domical volcanoes without postdating features (squares), domical with postdating features
(circles), volcanoes with a depressed central region (triangles), and volcanoes with depressed
region that has a large interior edifice (stars).
Figure 6. Profile of topography and estimate of temperature anomaly on the lower
subsurface shell for Ushas Mons. Temperatures are calculated from the surface density on the
lower shell for the two-layer inversion. The surface density, centered at 200 km depth, is
converted to temperature by assuming a perturbing layer of thickness 200 km and a coefficient of
28
thermal expansion of 3 x 10-5 K-1.
Figure 7. Same as Figure 4d (Atanua Mons), but with upper subsurface shell placed at a
depth of 15 km.
Figure 8. Same as Figure 4d (Atanua Mons), but inversion is only carried out to degree and
order 60, with a cos2 taper from degrees 40 to 60.
Figure 9. Spatio-spectral localizations for several volcanoes using Lwin = 15 (2500 km
window). Thick black line is admittance (mgal/km) and thin dotted line is gravity-topography
correlation. Thin lines are the admittance of the windowed topography compensated with elastic
lithospheric thicknesses of 10 km (solid lines) and 30 km (dashed lines), and crustal thicknesses
of 20 km (lower line of pair) and 40 km (upper line). Sif Mons, Hathor Mons, Unnamed (71.5
N, 256.0 E, and Uretsete Mons are type d in Table 1; Nagavonyi Corona is type ce; and Nzambi
Corona, Uti Hiata, and Mielikki Mons are type c.
Figure 10. Summary of spatio-spectral localizations with mean (solid line) and standard
deviations (dashed line) for domical volcanoes with no postdating features (thick line), domical
volcanoes with postdating features (medium line) and nondomical volcanoes (thin line).
29
Tables
Table 1. General properties of studied volcanoes.
Volcano Name
Ushas Mons
Sif Mons
Tepev Mons
Dzalarhons Mons
Idunn Mons
Unnamed
Innini Mons
Hathor Mons
Unnamed
Renpet Mons
Var Mons
Api Mons
Uretsete Mons
Furki Tholus
Kunapipi Mons
Tuulikki Mons
Nyx Mons
Chloris Mons
Nagavonyi Corona
Nefertiti
Kokyanwuti Mons
Nzambi Corona
Unnamed
Uti Hiata
Kunhild
Atanua Mons
Mielikki Mons
Atira Mons
Ituana Corona
Latitude Longitude Type
-25.0
323.5
d
22.0
352.0
d
29.5
45.0
d
0.0
34.0
d
-46.5
215.0
d
71.5
256.0
d
-34.5
328.5
d
-39.0
324.5
d
2.5
45.5
d
76.0
235.0
d
2.0
316.0
d
38.5
55.0
d
-12.5
261.5
d
35.0
236.0
d
-34.0
86.0
d
10.0
274.5
d
30.0
49.0 ce
-45.5
294.0 ce
-18.0
259.0 ce
36.5
48.0
c
35.5
211.5
c
-45.5
287.0
c
38.0
323.0
c
16.0
69.0
c
19.0
80.0
c
9.5
308.5
c
-28.0
281.0
c
52.5
267.5
c
19.5
154.0
c
Height (km)
2.0
2.5
5.0
3.0
3.0
2.3
2.3
2.3
1.8
2.0
2.0
1.8
2.3
1.5
2.5
2.0
1.3
0.6
0.8
2.1
1.3
0.7
1.0
1.8
1.5
1.5
1.5
1.3
0.8
Width (km)
Isostatic Anomaly
Half-Width 500-m Width Postdating mGal Separable
213
225
x
60
y
125
175
x
46
y
138
188
x
45
y
113
200
x
45
y
140
200
x
40
y
360
510
F,R
29
y
180
250
x
28
y
288
350
x
28
y
138
155
x
25
y
260
400
C
11
n
75
160
x
6
y
200
242
R
3
n
125
142
E,F
0
n
100
140
R,F
-10
n
200
308
C,F,R
-12
y
225
275
E,R
-26
n
400
500
E
65
y
140
x
R,F
16
n
200
175
F
-11
y
400
500
C
24
n
200
250
C,E,F
3
n
200
188
C,E,R,L
2
n
270
270
C,R,F
-4
n
263
288
E,F,R
-5
n
250
400
C,R
-10
n
275
275
x
-11
n
300
300
C,L,E,F
-12
n
190
200
E,R
-14
n
200
163
E,C,R
-18
n
Latitude is degrees north, longitude is degrees east. Type characterizes the shape as having a
broad domical shape with no significant topographic depression (d), containing a broad central
depression (c), or containing a broad central depression with volcanic edifices of significant size
in their interiors (ce). Height is given relative to the surrounding terrain. Width is given as both
the width at half of the maximum height of the volcano and at an elevation 500 m above the
surrounding terrain. Postdating features unrelated to the volcano are impact craters (C),
embayment by flows from another volcano (E), wrinkle ridges (R), fractures (F), and lineaments
(L) of unknown origin. Isostatic anomaly is calculated assuming a 30-km compensation depth
and a crustal density of 2900 kg m-3, and last column indicates whether the anomaly is clearly
separable and associated with the volcano.
30
Table 2. Dimensions of central depressed regions
Volcano Name
Nefertiti
Kokyanwuti Mons
Nzambi Corona
Unnamed (38N, 323E)
Uti Hiata
Kunhild
Atanua Mons
Mielikki Mons
Atira Mons
Ituana Corona
Depth (km)
0.9
0.7
0.4
0.3
0.4
0.5
0.4
0.3
0.3
0.8
Width (km)
250
125
175
125
125
150
300
175
100
125
Table 3. Summary statistics of isostatic anomalies
Domical (d), all
Domical without postdating features
Domical with postdating features
Depressed central region (c )
Edifice fills depression (ce)
Mean (mGal)
20
36
-1
-5
23
Std. deviation
25
16
18
12
39
Median
27
40
0
-8
16
31
Table 4. Summation of results for 2-layer inversions
Name
Ushas Mons
Sif Mons
Tepev Mons
Dzalarhons Mons
Idunn Mons
Unnamed (71.5N, 256E)
Innini Mons
Hathor Mons
Unnamed (2.5N, 45.5E)
Renpet Mons
Var Mons
Api Mons
Uretsete Mons
Furki Tholus
Kunapipi Mons
Tuulikki Mons
Nyx Mons
Chloris Mons
Nagavonyi Corona
Nefertiti
Kokyanwuti Mons
Nzambi Corona
Unnamed (38N, 323E)
Uti Hiata
Kunhild
Atanua Mons
Mielikki Mons
Atira Mons
Ituana Corona
Type
d
d
d
d
d
d
d
d
d
d
d
d
d
d
d
d
ce
ce
ce
c
c
c
c
c
c
c
c
c
c
Postdating
x
x
x
x
x
F,R
x
x
x
C
x
R
E,F
R,F
C,F,R
E,R
E
R,F
F
C
C,E,F
C,E,R,L
C,R,F
E,F,R
C,R
x
C,L,E
E,R
E,C,R
Upper Surface
10^6 kg/m^2 separable
0.5
n
0.5
n
0.7
n
0.6
n
-1.8
y
-3.9
y
-0.6
y
-2.1
y
-0.6
n
-1.7
y
-0.7
n
1.1
y
-1.9
y
-0.4
y
-3.4
y
-2.5
y
-1.4
y
-0.6
n
-2.4
y
-1.1
y
-0.4
y
-1.8
y
-0.6
y
-2.3
y
-3.2
y
-2.3
y
-3.8
y
-2.4
y
0.0
y
Lower Surface
10^6 kg/m^2 separable
-7.0
y
-5.1
y
-7.1
y
-5.5
y
-1.9
n
-1.7
y
-3.8
y
-4.1
y
-1.9
y
0.3
n
-0.1
y
-1.0
n
0.1
n
1.3
n
2.2
n
2.5
y
-5.9
y
-1.8
n
0.0
n
-2.8
n
-0.3
n
0.3
n
1.9
n
1.5
y
1.8
y
2.0
y
1.0
n
3.4
n
2.7
y
Volcanoes are ordered as in Table 1. Surface density values are located over the center of the
volcano in units of 106 kg m-2. Columns labeled "separable" indicate whether a surface density
anomaly is centered over the volcano and clearly separable from nearby features.
32
Table 5. Summation of spatio-spectral localizations
Name
Ushas Mons
Sif Mons
Tepev Mons
Dzalarhons Mons
Idunn Mons
Unnamed (71.5N, 256E)
Innini Mons
Hathor Mons
Unnamed (2.5N, 45.5E)
Renpet Mons
Var Mons
Api Mons
Uretsete Mons
Furki Tholus
Kunapipi Mons
Tuulikki Mons
Nyx Mons
Chloris Mons
Nagavonyi Corona
Nefertiti
Kokyanwuti Mons
Nzambi Corona
Unnamed (38N, 323E)
Uti Hiata
Kunhild
Atanua Mons
Mielikki Mons
Atira Mons
Ituana Corona
Type
d
d
d
d
d
d
d
d
d
d
d
d
d
d
d
d
ce
ce
ce
c
c
c
c
c
c
c
c
c
c
Postdating Degree 20
x
0.114
x
0.247
x
0.214
x
0.167
x
0.277
F,R
0.127
x
0.111
x
0.156
x
0.090
C
0.160
x
0.082
R
0.182
E,F
0.092
R,F
0.140
C,F,R
0.093
E,R
0.156
E
0.233
R,F
0.071
F
0.067
C
0.217
C,E,F
0.142
C,E,R,L
0.098
C,R,F
0.081
E,F,R
0.076
C,R
0.098
x
0.068
C,L,E
0.122
E,R
0.071
E,C,R
0.088
RMS gravity (mGal)
40
60
0.070
0.032
0.055
0.037
0.079
0.056
0.081
0.040
0.059
0.020
0.036
0.041
0.057
0.032
0.052
0.038
0.073
0.033
0.033
0.036
0.022
0.012
0.050
0.033
0.037
0.034
0.056
0.034
0.025
0.024
0.078
0.048
0.080
0.055
0.030
0.026
0.039
0.045
0.064
0.046
0.054
0.039
0.033
0.035
0.025
0.025
0.025
0.017
0.029
0.022
0.022
0.020
0.049
0.025
0.032
0.022
0.025
0.020
80
0.045
0.028
0.053
0.024
0.010
0.025
0.041
0.035
0.024
0.025
0.014
0.029
0.033
0.026
0.021
0.049
0.050
0.030
0.041
0.041
0.034
0.036
0.014
0.020
0.018
0.019
0.037
0.023
0.013
Admittance (mGal/km)
20
40
60
43
65
66
42
59
66
38
39
46
48
58
55
47
51
-1
32
20
39
34
43
48
39
40
54
25
37
34
35
23
39
30
32
6
28
31
31
34
32
36
42
45
47
33
13
33
38
52
56
38
37
45
29
21
32
25
29
42
35
34
36
37
47
49
28
27
38
29
31
55
24
10
17
26
15
21
31
23
1
38
32
22
21
24
33
39
26
35
80
94
59
90
45
11
33
52
42
35
27
36
44
46
41
41
80
82
41
63
57
49
49
13
33
32
28
54
49
38
20
0.99
0.98
0.97
0.96
1.00
0.91
0.94
0.97
0.73
0.95
0.95
0.81
0.94
0.96
0.93
0.92
0.97
0.96
0.93
0.90
0.98
0.98
0.93
0.61
0.83
0.97
0.91
0.84
0.98
Correlation
40
60
0.95
0.70
0.91
0.87
0.93
0.84
0.84
0.83
0.95
-0.02
0.90
0.80
0.90
0.78
0.91
0.82
0.81
0.75
0.84
0.76
0.82
0.11
0.81
0.78
0.81
0.56
0.86
0.79
0.50
0.85
0.94
0.69
0.92
0.83
0.63
0.74
0.88
0.62
0.87
0.82
0.92
0.79
0.77
0.80
0.77
0.56
0.58
0.75
0.68
0.74
0.67
0.01
0.77
0.58
0.65
0.53
0.58
0.56
80
0.84
0.83
0.86
0.82
0.36
0.76
0.80
0.79
0.63
0.66
0.56
0.81
0.66
0.65
0.68
0.77
0.83
0.67
0.74
0.76
0.69
0.72
0.22
0.80
0.75
0.46
0.65
0.58
0.65
33
Topography (km)
Figure 1.
3
2.5
2
1.5
1
0.5
0
0
100
200
300
400
500
600
500
600
Topography (km)
N- S profile (km)
3
2.5
2
1.5
1
0.5
0
0
100
200
300
400
W-E profile (km)
a. Sif Mons
34
Topography (km)
Fig 1 cont.
2.5
2
1.5
1
0.5
0
-0.5
0
100
200
300
400
500
600
500
600
Topography (km)
N- S profile (km)
2.5
2
1.5
1
0.5
0
-0.5
0
100
200
300
400
W-E profile (km)
b. Innini Mons
35
Topography (km)
Fig 1 cont.
2.5
2
1.5
1
0.5
0
-0.5
-1
0
100 200 300 400 500 600 700
Topography (km)
N- S profile (km)
2.5
2
1.5
1
0.5
0
-0.5
-1
0
100 200 300 400 500 600 700
W-E profile (km)
c. Kunapipi Mons
36
Topography (km)
Fig 1 cont.
1
0.5
0
0
100
200
300
400
300
400
Topography (km)
N- S profile (km)
1
0.5
0
0
100
200
W-E profile (km)
d. Chloris Mons
37
Topography (km)
Fig 1 cont.
1.5
1
0.5
0
-0.5
-1
0
100
200
300
400
500
400
500
Topography (km)
N- S profile (km)
1.5
1
0.5
0
-0.5
-1
0
100
200
300
W-E profile (km)
e. Kokyanwuti Mons
38
Topography (km)
Fig 1 cont.
1
0.5
0
-0.5
0
100
200
300
400
500
600
500
600
Topography (km)
N- S profile (km)
1
0.5
0
-0.5
0
100
200
300
400
W-E profile (km)
f. Atanua Mons
39
Figure 2.
Isostatic anom (mGals)
a.
60
50
40
30
0
100
200
300
400
500
600
500
600
Isostatic anom (mGals)
N-S profile (km)
60
50
40
30
0
100
200
300
400
W-E profile (km)
40
Fig 2 cont.
Isostatic anom (mGals)
b.
10
5
0
-5
-10
0
100
200
300
400
500
400
500
Isostatic anom (mGals)
N- S profile (km)
10
5
0
-5
-10
0
100
200
300
W-E profile (km)
41
Figure 3.
42
Figure 4.
a. Ushas Mons
43
Fig. 4 cont.
b. Hathor Mons and Innini Mons.
44
Fig. 4 cont.
c. Tuulikki Mons
45
Fig 4 cont.
d. Atanua Mons
46
Figure 5.
47
Figure 6.
48
Figure 7.
49
Figure 8.
50
Figure 9.
51
Fig. 9 cont.
52
Fig. 9 cont.
53
Fig. 9 cont.
54
Figure 10.
55
Fig. 10 cont.
56