4574 - Ivanov, M. A., and J. W. Head III

Transcription

Planetary and Space Science 113-114 (2015) 10–32
Contents lists available at ScienceDirect
Planetary and Space Science
journal homepage: www.elsevier.com/locate/pss
The history of tectonism on Venus: A stratigraphic analysis
Mikhail A. Ivanov a,b,n, James W. Head b
a
b
V. I. Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences, 19 Kosygin Street, 119991 Moscow, Russia
Department of Earth, Environmental and Planetary Sciences, Providence, RI 02912, USA
art ic l e i nf o
a b s t r a c t
Article history:
Received 2 June 2014
Received in revised form
23 January 2015
Accepted 20 March 2015
Available online 30 March 2015
The surface of Venus displays several tectonized terrains in which the morphologic characteristics of the
original materials are almost completely erased by superposed tectonic structures whose large dimensions (»100 km) suggest formation related to mantle convection. The characteristics of these tectonized
terrains are in contrast to volcanic units in which tectonic structures are less significant or absent and
thus do not obscure the volcanic character of the units. We describe the temporal distribution of tectonized terrains, their stratigraphic relationships with volcanic units, and how these outline the major
episodes in the geological evolution of Venus. Five major tectonized units make up 20% of the planet:
(1) tessera (t, 7.3%), (2) densely lineated plains (pdl, 1.6%), (3) ridged plains/ridge belts (pr/rb, 2.4%),
(4) groove belts (gb, 8.1%), and (5) rift zones (rz, 5.0%). Clear relationships of relative age are often seen
among the tectonic and volcanic units at the global scale and define three contrasting regimes of volcanic
and tectonic resurfacing. The majority of tectonized terrains (t through gb) are the products of tectonic
resurfacing and are embayed by the vast volcanic plains and, thus, are older. There are no units with
either mildly- or non-tectonized surfaces that interleave the tectonic terrains, which would be expected
if the tectonic resurfacing operated only during specific repetitive phases in discrete regions. These
tectonized terrains (t through gb) thus define a tectonically dominated regime of resurfacing that occurred
at a global-scale near the beginning of the observable geological history of Venus. This ancient tectonic
regime began with formation of tessera and was followed by formation of pdl and pr/rb. Groove belts
formed near the end of this regime. Branches of groove belts compose the tectonic components of many
coronae, suggesting that these features are genetically related (e.g., mutual development of mantle
diapirs and zones of extension) and that coronae may have punctuated the final stages of the ancient
tectonic regime. This regime was followed by emplacement of the vast volcanic plains, such as shield and
regional plains, the surfaces of which are extensively deformed by the global network of wrinkle ridges.
Emplacement of the plains defines the second, volcanically dominated regime, representing a time when
surface tectonic deformation related to the mantle convection waned. Rift zones are the stratigraphically
youngest manifestations of regional-scale tectonic deformation on Venus. Rifts are spatially and temporarily associated with the youngest lava flows and often cut the crest areas of large, but isolated, domeshaped rises. Structures of rift zones always cut the surface of the vast plains, which means that rifts are
separated in time from the ancient tectonic regime, post-date the regional plains, and represent a new
phase of tectonism that was contemporaneous with the late volcanism of lobate plains. Rift zones and
lobate plains define the third, network rifting-volcanism regime, of resurfacing that was related to late
stages of evolution of the dome-shaped rises.
& 2015 Elsevier Ltd. All rights reserved.
Keywords:
Venus
Tectonism
Tectonic evolution
Resurfacing regimes
1. Introduction
Regional-scale tectonic structures are the most prominent manifestations of the loss of internal heat of a planet (Solomon and Head,
1982). This is the case for features of different extents and abundances, such as structural assemblages related to plate tectonics on
n
Corresponding author.
E-mail address: [email protected] (M.A. Ivanov).
http://dx.doi.org/10.1016/j.pss.2015.03.016
0032-0633/& 2015 Elsevier Ltd. All rights reserved.
Earth, zones of bright terrains on Ganymede (Collins et al., 2013),
and even suites of mare-type wrinkle ridges on volcanic plains of
Mars and the Moon. Since the Venera-15/16 and Magellan missions,
it is a well-established fact that Venus possesses a rich variety of
tectonic landforms (Barsukov et al., 1986; Solomon et al., 1992),
which distinguishes Venus from the smaller terrestrial planetary
bodies (e.g., Mars, Mercury and the Moon) and makes it more
similar to the Earth.
Because Venus is almost as large as Earth and has a similar bulk
density (hence, bulk composition) (Barsukov, 1992), it has both
M.A. Ivanov, J.W. Head / Planetary and Space Science 113-114 (2015) 10–32
gravitational and compositional potential for long-lasting tectonic
activity that continued at least for 85–90% of the total geologic
history of the planet (Arkani-Hamed and Toksoz, 1984; Parmentier
and Hess, 1992). These are manifested by numerous tectonic structures, some of which characterize relatively recent (tens of million
years ago) episodes of the visible portion of the geologic history of the
planet (e.g., Basilevsky and Head, 2000; McGill et al., 2012).
Atmospheric conditions on Venus (Moroz, 1983; Seif, 1983;
Avduevsky et al., 1983) preclude the existence of liquid water near
the surface and suppress wind activity causing the lack of substantial
erosional/sedimentation processes on the surface of the planet
(Arvidson et al., 1991, 1992). The number of impact craters on Venus
is small, 1000 (Herrick et al., 1997; Schaber et al., 1998). These
factors result in the fact that tectonic and volcanic landforms are the
principal contributors to the surface geology (Barsukov et al., 1986;
Head et al., 1992; Solomon et al., 1992). The difference in the nature
of the principal features that define terrains of either tectonic or
volcanic origin indicates that they reflect different styles of resurfacing: emplacement of volcanic units due to eruption and accumulation of volcanic materials and deformation and displacement of
these materials during their re-working by tectonism.
A global geological map of Venus at the 1:10 M scale (Ivanov
and Head, 2011) shows the spatial distribution and age relationships of the volcanic and tectonic terrains and provides the possibility to address several important problems related to the geologic evolution of the planet:
1. What is the variety of volcanic and tectonic landforms on
Venus and what is their stratigraphic position and their relationships? What is the nature of the processes that are manifested by specific landforms?
2. What style of tectonic activity and its evolution are recorded in
the observable suites of tectonic structures/terrains? What
tectonic features may indicate the prevalence of either horizontal or vertical displacements of crust/lithosphere? How did
the regional/global pattern of stresses on Venus evolve as a
function of time?
3. What do the typical relationships of relative ages among the
tectonized terrains and between them and volcanic units tell us
about the structure of the visible portion of geologic history of
Venus?
The spatial and temporal distributions of the volcanic units and
the observable history of volcanism on Venus have been described
in a separate paper (Ivanov and Head, 2013). In this paper we
address the problem of the evolution of tectonic activity of the
planet based on a review of tectonic landforms and their stratigraphic position as they are seen globally.
2. Methods
Our work is based on the analysis of the global geological map
of Venus (1:10 M scale, Ivanov and Head, 2011) that portrays the
spatial and temporal distribution of landforms of different origin.
The map is a result of a three-fold procedure, the first of which is
the definition of units. We used the traditional way to define units:
If the surface of a terrain has a specific and distinctive morphology,
which is different from the adjacent/remote terrains, this morphology defines a mappable unit (e.g., Wilhelms, 1974).
On Venus, both volcanic and tectonic features commonly occur
together and characterize local and regional morphologies. The
abundance and scale of the features of either volcanic or tectonic
origin, however, vary significantly and define two major groups of
terrains: volcanic and tectonic units. Volcanic units (Ivanov and
Head, 2011, 2013) are those in which the original morphology
11
related to the emplacement of volcanic materials is obvious and
tectonic structures are subordinate and do not erase the initial
morphology of a unit. For example, regional plains of Venus in
many respects resemble the Hesperian ridged plains on Mars
(Greeley and Guest, 1987). On both planets, the original surface of
the plains is subsequently deformed by numerous wrinkle ridges
but retains its pre-deformational morphology.
The situation is complicated if tectonic structures strongly
affect the surface of a terrain. The structural component can be so
pervasive that it almost completely obscures/erases the morphological characteristics of the underlying materials. In these cases,
the structure becomes a part of the definition of terrains that
characterize the surface of Venus. A good example of such a unit
on Venus is tessera that certainly consists of some materials but
their original characteristics are systematically overprinted by
numerous secondary tectonic structures (e.g., Solomon et al., 1992;
Ivanov and Head, 1996; Hansen and Willis, 1996). Terrains of this
type represent material-structural units for which it is impossible
to reliably separate the material and tectonic components at the
given resolution of images and scale of mapping. Tessera, densely
lineated plains, ridged plains, and mountain belts of Venus are
examples of such units, which are portrayed in many regional
geological maps of Venus (e.g., Rosenberg and McGill, 2001;
Campbell and Clark, 2006; Copp and Guest, 2007; Ivanov and
Head, 2010). These tectonized units are in contrast to the networks
of wrinkle ridges. Although the ridges are very abundant and
pervasive (Bilotti and Suppe, 1995), they do not modify underlying
materials beyond recognition and, thus, do not define specific
material-structural units.
Swarms of densely packed extensional structures are prominent
features of the surface of Venus. Usually, graben and fractures are
mapped as individual features independently of the material units on
which they are superposed. In some places, however, the density of
fractures is so great that they almost completely obscure the characteristics of underlying materials at the scale of the mapping. When
this occurs, the guidelines formulated by Wilhelms (1990) can be
applied: “when the deformed rock units are not recognizable … it is
better to map such structural units as ‘fractured plains material’ …
than to ignore the presence of the structures in order to adhere
strictly to the Code” (Wilhelms, 1990, pp. 227–228). Groove belts and
rift zones represent examples of these terrains. They are defined as
purely structural units, the tectonic elements of which destroy a
variety of preexisting materials. Again, the density of structures and
scale of mapping define the way of portrayal of these terrains either
as specific structural units (e.g., Basilevsky, 2008) or as swarms of
linear features (e.g., Grosfils et al., 2011).
The next step in the compilation of a geological map of the
local, regional, or global scale is portrayal of the defined units.
During this procedure, the characteristics of the contacts between
the units are considered and mapped. The result of this step is a
geomorphic map that shows the spatial distribution of units.
Finally, during the third step of the compilation of a geological
map the relationships of embayment and crosscutting between
the mapped units are analyzed. This provides the information on
the relative age relationships among the units. During this step,
the geomorphic map is converted into a geological map that shows
the distribution of units both in space and time.
On planets with a rich and diverse crater record, the crater
density is a standard means of validation of the relative ages
among the mapped units. On Venus, with its small number of
impact craters, this method is not applicable in most cases and the
relative age relationships play a key role in the assessment of the
local stratigraphy. The time correlation of the remote occurrences
of units with similar morphology requires mapping of extensive
contiguous regions that connect these occurrences (Ivanov and
Head, 2001a, 2001b). In such mapping, the most important are
12
reference units that have consistent age relationships with the
surrounding terrains. On Venus, the lower unit of regional plains
has the characteristics of a reference unit because (1) the plains
are extensive and pervasive; (2) they show about the same morphology in all areas where they occur and do not show features
that may suggest that the plains have a time-transgressive nature
(Ivanov and Head, 2011); (3) long lava channels, which can reach
thousands of kilometers in length, carve the surface of the plains
and suggest that the plains represent a single material unit; (4) the
plains demonstrate the consistent relationships of embayment and crosscutting with the surrounding units (Ivanov and
Head, 2011). Documentation of this unit during mapping in individual regions (Ivanov and Head, 2004a, 2005a, 2005b, 2008a,
2008b) and in geotraverses at 30°N (Ivanov and Head, 2001a,
2001b) and along the equator (Ivanov and Head, 2006) has proven
that the plains indeed represent the reference unit that can be
used for correlation of remote regions (Ivanov and Head, 2011).
The global geological map of Venus reveals that five tectonic
and four volcanic units make up about 96% of the surface of the
planet (Table 1). These units are the major contributors to resurfacing and, thus, are most important for the understanding of the
visible portion of the geologic history of Venus.
The group of tectonic units (Fig. 1) consists of tessera (t), densely
lineated plains (pdl), ridged plains (pr) that often make ridge belts
(RB), groove belts (gb), and rift zones (rz). In these units, which make
up about a quarter of the surface of Venus (Table 1), tectonic structures override the original morphology of underlying terrain and
dominate the appearance. In sharp contrast to the tectonic units, the
surfaces of volcanic units are usually morphologically smooth and are
cut by sparse tectonic structures that do not change significantly the
original morphology of volcanic materials (Fig. 1). The group of major
volcanic units makes up about 70% of the surface of Venus (Table 1)
and consists of shield plains (psh), the lower (rp1) and upper (rp2)
units of regional plains, and lobate plains (pl).
of kilometers. For example, the smaller Alpha–Lada cluster (A–L in
Fig. 2) is about 2000 9000 km2 and the largest cluster, Ovda–Thetis
(O–T in Fig. 2) is about 5000 14,000 km2 across.
3. Definition and description of major tectonic units/terrains
on Venus
3.1. Tessera terrain (unit t)
Tessera (Fig. 1a) is one of the most tectonically deformed types of
terrain on Venus (e.g., Barsukov et al., 1986) and occupies about 8% of
its surface (Table 1). Tessera massifs are mostly equidimensional or
slightly elongated and vary in size from a few tens up to a few
thousands of kilometers (Ivanov and Head, 1996). Tesserae tend to
occur within several diffuse clusters (Fig. 2) extending for thousands
Table 1
Areas of the major tectonic and volcanic units on Venus.
Unit
Unit area, 106 km2
Unit area, %
Major tectonic units
t
pdl
pr
gb
rz
Sum
35.7
7.8
10.3
39.9
24.3
117.9
7.8
1.7
2.2
8.7
5.3
25.6
Major volcanic units
psh
rp1
rp2
pl
Sum
Total
85.2
152.0
45.2
40.7
323.0
440.9
18.5
33.0
9.8
8.8
70.2
95.8
Fig. 1. The groups of tectonic and volcanic units on Venus. Tectonic units: t –
tessera, pdl – densely lineated plains, pr/Rb – ridged plains/ridge belts, gb – groove
belts, rz – rift zones. Volcanic units: psh – shield plains, rp1 – the lower subunit of
regional plains, rp2 – the upper subunit of regional plains, pl – lobate plains. Both
tectonic and volcanic units are arranged in a stratigraphic order from older (top) to
younger (bottom) units. See text for explanation.
13
Fig. 2. Spatial distribution of tessera terrain (red) on Venus. The background is the topographic map in simple cylindrical projection. Capital letters indicate the tessera
clusters. I–A: Ishtar–Ananke; O–T: Ovda–Thetis; V–T: Virilis–Themis; A–L: Alpha Lada; N: Nemesis. Yellow dashed line show approximate contours of the clusters. The
thicker black solid line corresponds to 0 km contour (6051 km), the thinner black solid line-contour þ 1 km, and the thinner black dashed line-contour 1 km. (For
interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
The tessera surface always stands topographically higher than
the surroundings and tessera massifs are concentrated in regions
that lie above the zero contour line (Fig. 2). The largest tesserae
(e.g., Ovda, Thetis) correspond to a specific, plateau-like, class of
regional highlands on Venus (Sjogren et al., 1983). The tessera
hypsogram is shifted to the right (Fig. 3) and the largest tessera
regions make the second, high-elevation, peak of the hypsogram.
At least two sets of intersecting contractional (ridges) and
extensional (graben, fractures, scarps) structures characterize tessera
(Bindschadler and Head, 1991; Sukhanov, 1992). These structures
occur at broadly varying scales. The smaller structures (several kilometers wide, tens of kilometers long) are densely packed and often
erase the morphological characteristics of the tessera precursor
material (Fig. 1). Among the largest structures are relatively narrow
(a few tens of kilometers wide) but very long (hundreds of kilometers) troughs that divide the surface of tessera regions into individual large (hundreds of kilometers across) segments. Intratessera
lava plains often partly cover the floor of the troughs.
The age relationships between the contractional and extensional structures are of extreme importance for the assessment of
the models of tessera formation (Ivanov and Head, 1996; Hansen
and Willis, 1996; Tuckwell and Ghail, 2003; Kumar, 2005) but are
usually unclear and in many cases cannot be established unambiguously because of the lack of stratigraphic markers. Impact
craters and/or intratessera plains may represent such markers
within tessera terrains. The survey of craters superposed on tessera has shown that tessera ridges deformed none of them, but
some of the craters are cut by the extensional structures (Gilmore
et al., 1997, 1998). Although these relationships suggest that the
ridges are older, the number of craters in tessera is small and only
in rare cases can they serve as local stratigraphic markers.
The intratessera plains are much more abundant and provide
more useful markers that help to determine the relative ages of
the tessera ridges and graben (Ivanov and Head, 2000). Usually,
the plains postdate both sets of tessera structures but in some
regions they intervene between the ridges and graben and thus
provide a robust indicator of the stratigraphic relationships of the
tessera structures.
In the central portion of Ovda Tessera (Fig. 4) the surface
consists of low NW-trending ridges with a fine-scale lineation (r1
and r2) that are surrounded by intratessera plains (pl). The surface
of the plains is flat, which suggests that they have been emplaced
after the episodes of formation of ridges and, thus, are younger. A
series of NE-trending low scarps represents the third type of features in this area. The scarps are often paired and inward looking
and form narrow and shallow graben that in this region were
called “ribbons” by Hansen et al. (2000). The scarps/graben cut
both the tessera ridges and the plains and, thus, are younger.
In the eastern portion of Fortuna Tessera (Fig. 5) the surface
consists of densely packed ridges of NW orientation and numerous
NNE-trending narrow grooves that are morphologically similar to
the “ribbon”-type graben in the western portion of Fortuna
(Pritchard et al., 1997; Hansen and Willis, 1998). The relationships
between these structures are often ambiguous: they could have
been formed either simultaneously, or the graben may be older, or
vice versa. A large ( 50 75 km2) occurrence of intratessera
plains in this area serves as a local stratigraphic marker that provides the means to resolve this problem: materials of the plains
embay the tessera ridges and are cut by the graben indicating the
older age of the ridges.
In many places within Ovda Tessera, at least two generations of
the intratessera plains are presented (Fig. 6). The younger plains
(plit2) are featureless, tectonically undeformed, and their surface
appears to be flat and nearly horizontal. The older plains (plit1)
embay flanks of the NE-trending tessera ridges and are tilted away
from them. Series of narrow NW-trending graben (“ribbons”, Ghent
and Hansen (1999)) cut both the tessera ridges and the older intratessera plains and are embayed by the younger, sub-horizontal plains
(Fig. 6). These relationships unambiguously suggest that the contractional structures (ridges) are the oldest in this region.
In all occurrences of tessera terrain, we have conducted a systematic search of areas where the age relationships between the
contractional and extensional structures can be established with
the help of local stratigraphic markers. These areas have been
found in the eastern portion of Fortuna Tessera, throughout the
largest tesserae of Ovda and Thetis Regiones, and in tessera to the
north of Thetis (Haasttse-baad Tessera). In all these regions (Fig. 7)
14
Fig. 3. Hypsograms of the major tectonic and volcanic units on Venus (thick line) in comparison with the global hypsogram (thin line). The lowland, midland, and highland
topographic domains are defined in Masursky et al. (1980). See (Ivanov and Head, 2013, Table 4) for the statistics of the unit topographic distribution.
evidence for the older age of graben is absent and the stratigraphic
markers of the intratessera plains show relationships similar to
those shown in Figs. 4–6 that indicate the older age of the ridges.
3.2. Densely lineated plains (unit pdl)
The defining characteristic of this unit is that its surface is
heavily dissected by numerous densely packed linear and curvilinear lineaments (Fig. 1b). The lineaments (fractures) are narrow
(few hundred meters wide), short (a few tens of kilometers long),
and parallel or subparallel to each other (Basilevsky and Head,
1995a, 1995b; Ivanov and Head, 2001a, 2011); the mean spacing of
the lineaments is 1.571.2 km (Mastrapa, 1997). The tectonic
structures of densely lineated plains almost completely erase the
original morphologic characteristics of the underlying materials
(Fig. 1b).
Densely lineated plains usually occur as small (tens to a few
hundreds of kilometers across) outcrops, the surface of which is
above the surrounding volcanic plains. The main peak of the
hypsogram of unit pdl is slightly shifted toward the higher elevation relative to the total hypsometric curve (Fig. 3).
Although the total area of densely lineated plains is small ( 2% of
the surface of Venus, Table 1) pieces of the plains are more evenly
distributed over the surface of Venus than those of tessera and
15
Fig. 4. An example of stratigraphic relationships of contractional and extensional structures in tessera where lava plains postdate tessera ridges but predate narrow graben
(“ribbons”). The figure shows a piece of the surface in the central portion of Ovda tessera. Part of F-MIDR 00S076, center of the image is at 4.3°S, 76.8°E.
Fig. 5. An example showing the age relationships between tessera ridges, graben, and intratessera plains. The plains embay tessera ridges but are cut by the graben. Part of
C1-MIDR 75N074, center of the image is at 69.3°N, 67.8°E.
16
Fig. 6. An example of stratigraphic relationships of contractional and extensional structures in tessera where lava plains serve as a local stratigraphic marker that indicates
the relative age of ridges and graben. The figure shows a piece of the surface in the northern-central portion of Ovda tessera. The older intratessera plains (plit1) embay the
tessera ridges and are cut by sets of narrow graben. Part of F-MIDR 00N082, center of the image is at 2.3°S, 90.0°E.
sometimes occur in groups where orientation of the lineaments is
approximately the same. The groups of the pdl occurrences are a few
thousands of kilometers across and tend to be at the periphery of the
tessera clusters but usually away from the tessera massifs (Fig. 8).
Only in several areas are densely lineated plains in contact with
tessera. In these localities, two situations are observed. More often,
material components of pdl embay tessera and their tectonic
structures cut tessera (Fig. 9a) (Basilevsky and Head, 1995a,
1995b). However, in some places the surface of densely lineated
plains is additionally deformed and the structural pattern of the
unit resembles that of tessera (Fig. 9b). This is so-called tessera
transitional terrain, ttt, (Ivanov and Head, 2001a).
For example, a large area of densely lineated plains to the west
of Lakshmi Planum (Atropos Tessera) in which the typical pdllineaments are oriented in a N direction, is additionally deformed
by sets of broad WNW-trending ridges. Intersections of these
structures produce a tessera-like structural pattern (Fig. 9b) (Ivanov and Head, 2008b). Along the north margin of Meskhent Tessera, occurrences of densely lineated plains are deformed by
structures that appear to continue the structural trend of the later
extensional structures of the tessera. These examples suggest that
although the plains are likely to postdate tessera, some portion of
the plains was involved in the latest phases of tessera deformation
(Ivanov and Head, 2008a) and both units (tessera and densely
lineated plains) are partly overlapping each other in time.
3.3. Ridged plains and ridge belts (unit pr/RB)
Linear and curvilinear ridges (5–10 km wide, several tens of
kilometers long) dominate the surface of ridged plains (Fig. 1c). A
characteristic feature of ridged plains is that their ridges are often
collected into prominent belts (ridge belts) (Frank and Head, 1990;
Kruychkov, 1992). Both the higher radar backscatter cross-section of
the ridged plains and the belts (McGill and Campbell, 2006) and the
specific structural pattern distinguish this unit from the other terrains on Venus. Ridged plains and ridge belts occupy about 2.2% of
the surface of Venus (Table 1) and usually form elongated occurrences hundreds of kilometers long and many tens (to a few hundred) kilometers wide (Fig. 10). The most abundant occurrences of
ridge belts are located within the northern hemisphere in a broad
17
Fig. 7. Examples of tessera where intratessera plains help to establish the relative ages of the tessera ridges and graben (arrows). In all cases, the plains postdate the ridges
and predate the graben. (a) Northwestern portion of tessera in Ovda Regio. Part of C1-MIDR 00N060, center of the image is at 5.7°N, 63.6°E. (b) Southern portion of tessera in
Ovda Regio. Part of C1-MIDR 15S077, center of the image is at 9.4°S, 80.3°E. (c) Portion of Tessera Haasttse-baad to the north of Thetis Regio. Part of C1-MIDR 00N129, center
of the image is at 4.1°N, 128.2°E. (d) Portion of tessera in Thetis Regio. Part of C1-MIDR 00N129, center of the image is at 3.3°S, 131.9°E. Scale bars are 25 km in each image.
fan-shaped area centered at 200°E (Barsukov et al., 1986; Frank
and Head, 1990; Kryuchkov, 1990, 1992) where the belts are concentrated at the edges of the Nemesis (N) and Virilis–Themis (V–T)
tessera clusters (Fig. 10). Prominent ridge belts characterize the
northern flank of the Ovda–Thetis cluster (O–T) and are associated
with large tessera massifs within the Ishtar–Ananke tessera cluster
(I–A). A group of smaller occurrences of the belts populates the floor
of Lavinia Planitia (Squyres et al., 1992; Koenig and Aydin, 1998;
Ivanov and Head, 2001b; Fernandez et al., 2010) to the west of the
Alpha–Lada tessera cluster (A–L). Thus, the belts preferentially occur
either at the periphery of the tessera clusters or at the edges of the
large tessera massifs.
Both ridged plains and ridge belts have higher relief than the
surrounding plains units and the hypsogram of ridged plains is
noticeably shifted toward higher elevations relative to the total
hypsogram (Fig. 3). This suggests that the plains/belts are less sensitive to later embayment and flooding than the patches of densely
lineated plains and their actual distribution is likely to mark the
regions where the plains were more intensively deformed.
In this research, we consider mountain belts (unit mb (Ivanov
and Head, 2011)) that surround Lakshmi Planum (Barsukov et al.,
1986; Head, 1990; Pronin, 1992) (Fig. 10) as a topographic variety
of ridge belts. Mountain belts resemble ridge belts morphologically but strongly differ from them in their topographic characteristics and represent the only real mountain ranges on Venus
(Pettengill et al., 1980; Masursky et al., 1980); they make up a
negligible portion of the surface of Venus ( 0.3%).
In places where ridged plains occur in contact with tessera,
material of the ridged plains usually embays the tessera (Fig. 11a,
white arrow) and ridges of ridged plains cut the tessera surface
(Fig. 11a, black arrow). This means that both emplacement of material of ridged plains and its deformation postdated the main phases of
tessera formation. Many of the fragments of ridged plains, however,
are away from massifs of tessera (Fig. 10) and it precludes direct and
reliable determination of the relative ages between these units in
18
Fig. 8. Spatial distribution of densely lineated plains (red) on Venus. Tessera massifs (light gray) and boundaries of the tessera clusters are shown for reference. The background is
the topographic map in simple cylindrical projection. The thicker black solid line corresponds to 0 km contour (6051 km), the thinner black solid line-contour þ1 km, and the
thinner black dashed line-contour 1 km. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
many cases. In some areas where ridged plains are in contact with
tessera, the transitional morphologies (the tessera transitional terrain, ttt) appear. For example, at the southern margin of Tellus Tessera, ridged plains embay the contractional structures of tessera, but
are deformed by sets of parallel grooves/graben that represent the
later stages of tessera deformation (Ivanov and Head, 2001a; Gilmore
et al., 2010; Graupner Bergmann and Hansen, 2013). The intersections between the ridges of ridged plains and the grooves result in
formation of a tessera-like structural pattern of the plains.
One of the most prominent examples of the tessera transitional
terrain, which is formed at the expense of ridged plains, is
observed at the northern edge of Ovda Regio (Fig. 11b). In this area,
a ridge belt is approaching the edge of tessera from the NW and
merges with it at 7°N, 95°E. To the east of this point, the belt
outlines the flank of tessera and the ridges of the belt that are
closer to tessera are cut by sets of graben that extend from the
tessera region (Fig. 11b, black arrows). At the distal portion of the
belt, the ridges are not deformed by the graben (Fig. 11b, white
arrow). This situation is observed within a relatively short segment
of the ridge belt (between 95 and 100°E). Further to the east, the
belts that are in contact with tessera of either Ovda or Thetis
Regiones do not show evidence for the tessera transitional terrain.
This means that the transitional morphologies were formed only
occasionally and suggests that the formation of tessera and ridge
belts was partly overlapping in time.
Pieces of densely lineated and ridged plains usually occur away
from each other (Figs. 8 and 10) but when they are in contact
fragments of pdl appear to be embayed by materials of ridged
plains (e.g., Ivanov and Head, 2008a). These relationships point to
the relatively older age of densely lineated plains, although formation of these units may have overlapped in time.
The mean spacing of the structures of groove belts is 3.172.5 km
(Mastrapa, 1997). The density of fractures/graben in groove belts is
usually so high that they almost completely obscure the characteristics of underlying materials at the scale of the mapping. Groove
belts represent a purely tectonic unit during the formation of which
various materials have been cut by fractures and graben.
Groove belts occupy 8.7% of the surface of Venus (Table 1) and
occur as zones of many hundreds (up to a few thousands) of kilometers long and tens to a few hundred kilometers wide (Fig. 12). In
contrast to ridge belts, groove belts do not show the spatial correlation with the tessera clusters, are distributed more broadly, and
form an anastomosing pattern of branches that divide the surface
into blocks hundreds to thousands of kilometers across (Fig. 12).
Inside the branches of the belts, their fractures and graben often
form elliptical and circular features, the floor of which is covered by
later plains materials. Individual branches of the belts very often are
curved and bifurcated and produce the circular rims of coronae
(Fig. 13). Because of this, coronae are often spatially associated with
groove belts (Fig. 14). Branches of groove belts always represent local
highs, the surface of which stands several hundred meters above the
surrounding plains and the majority of groove belts occur in regions
that are higher than the zero contour line (Fig. 12). The hypsogram of
the belts is noticeably shifted toward higher elevations relative to the
total hypsometric curve (Fig. 3).
In regions where groove belts are in contact with tessera,
densely lineated, and ridged plains the belts cut these units indicating the younger age of groove belts (Fig. 15). The large number
of groove belts that occur in many areas on Venus (Fig. 12) permit
a confident establishment of the stratigraphic position of groove
belts relative to all other heavily tectonized units.
3.5. Rift zones (unit rz)
3.4. Groove belts (unit GB)
Groove belts (Fig. 1d) have been defined as zones of densely
packed extensional structures (fractures and graben) (Basilevsky and
Head, 1998; Ivanov and Head, 2011). The typical widths of these
structures are about several hundred meters and up to 1–2 km, and
individual fractures can reach several tens of kilometers in length.
Rift zones (Fig. 1e), as well as groove belts, represent a pure
structural unit that consists of numerous and densely packed
extensional structures, fractures and graben. In many rift zones
there are deep (several kilometers) and steep-sided canyons or
valleys that can be several tens of kilometers wide. In contrast to
groove belts, structures of rift zones on average are broader,
longer, and somewhat less densely packed than structures of
groove belts; the characteristic spacing of structures in rift zones is
3.2 72.5 km (Mastrapa, 1997). Fractures and graben of rift zones
are either parallel to or crisscrossing each other and often have a
zigzag-like planform, which is not typical of structures of groove
belts (Fig. 16).
Rifts appear as broad (hundreds of kilometers wide (e.g.,
Guseva, 2008) and very prominent zones that extend for thousands of kilometers. Rift zones occupy 5.0% of the surface of
Venus (Table 1) and preferentially occur within the equatorial area
of Venus where they are associated with large highlands, such as
eastern Aphrodite, and outline the Beta–Atla–Themis (BAT) region
(Fig. 17). Although the major rift valleys usually represent deep
topographic depressions (chasmata), the close association of rifts
with regional highs causes a strong shift of the unit hypsogram
toward higher elevations (Fig. 3). We will consider the stratigraphic position of rift zones in the following sections of the paper.
Fig. 9. The age relationships of densely lineated plains (pdl) and tessera (t).
(a) Materials of the densely dissected plains (pdl) embay a piece of tessera (lower
left); lineaments of pdl penetrate into tessera (arrow). These relationships indicate
that both emplacement and deformation of pdl postdate tessera. Part of C1-MIDR
30N135, center of the image is at 31.9°N, 128.0°E. (b) In several regions (e.g.,
Atropos Tessera), densely lineated plains are additionally deformed by broad ridges.
This superposition of lineaments of pdl and the ridges produces the tessera-like
pattern of deformation (tessera transitional terrain, ttt). These relationships suggest
that formation of both tessera and densely lineated plains were partly overlapping.
Part of C1-MIDR 75N299, center of the image is at 68.7°N, 310.2°E.
19
4. Major volcanic units
Four major volcanic units, the surface of which is mildly
deformed by tectonic structures (Fig. 1), make up about 70% of the
surface of Venus (Table 1) (Ivanov and Head, 2011, 2013). These
units were described in detail elsewhere (Ivanov and Head, 2011,
2013) and in this paper we are summarizing their most characteristic features.
Shield plains (unit psh) have a great number of small (from
1–2 km up to 10 km) shield- and cone-like mounds (Fig. 1f) that
are interpreted as volcanic edifices (Aubele and Slyuta, 1990; Head
et al., 1992; Guest et al., 1992). The volcanic emplacement of
materials played the major role in formation of shield plains and
later tectonic deformation (mostly by wrinkle ridges) was much
less important. Shield plains are quite abundant on Venus and
their exposed area comprises 18.5% of the surface of Venus
(Table 1). The age relationships between individual shields and the
surrounding materials are often ambiguous (Addington, 2001) and
suggest that formation of the shields and intershield plains were
overlapping in time. In contrast, the stratigraphic relationships of
the unit of shield plains (both the shields and intershield plains)
with the surrounding plains units are more clear and suggest that
shield plains pre-date emplacement of regional plains (Ivanov and
Head, 2004b).
The lower sub-unit of regional plains (unit rp1) has a morphologically smooth surface with mostly homogeneous (mottled
locally), and relatively low, radar backscatter (Fig. 1g). Although
the regional networks of wrinkle ridges deform the surface of the
unit, the ridges do not obscure its original morphological characteristics. The lower sub-unit of regional plains is the most
abundant and ubiquitous unit on Venus. Its exposed occurrences
make up 33.0% of the surface of Venus (Table 1) and they can be
traced almost continuously around the globe.
The upper sub-unit of regional plains (unit rp2) is characterized by a morphologically smooth surface that is moderately
deformed by the same family of wrinkle ridges that cut the surface
of rp1 (Fig. 1h). In contrast to the relatively low albedo of unit rp1,
the albedo of the upper sub-unit of the plains is noticeably higher
(Fig. 1h) and the albedo pattern is either homogeneous or consists
of faint flow-like features. The upper sub-unit of regional plains is
less widespread and covers 9.8% of the surface of Venus (Table 1)
and its flow-like features superpose the surface of the lower
member of regional plains.
Lobate plains (unit pl) usually have morphologically smooth
surfaces that are occasionally disturbed by a few extensional features (fractures, graben, Fig. 1i). The radar albedo of lobate plains
consists of numerous bright and dark features that are interpreted
as individual lava flows (e.g., Head et al., 1992; Crumpler and
Aubele, 2000). Lobate plains are typically associated with large
volcanoes and often occur within the major dome-shaped rises
such as Beta and Atla Regiones, and Lada Terra. The flows of lobate
plains embay wrinkle ridges in all localities where the plains and
the ridges occur together. Thus, among the other major volcanic
units, lobate plains are the youngest (e.g., Campbell and Campbell,
2002; Stofan and Guest, 2003; McGill, 2004; Grosfils et al., 2011;
Ivanov and Head, 2011) and occupy 8.8% of the surface of Venus
(Table 1). Because of the stratigraphic position of the plains, their
exposed area is close to the true area of the unit.
5. Age relationships of tectonic terrains with volcanic units
Massifs of tessera show two types of boundaries with the surrounding volcanic units (Ivanov and Head, 1996): (1) very sinuous
contacts, which are characterized by deep penetrations of lava
plains into tessera, and (2) contacts that are more smooth at the
20
scale of hundreds of kilometers but show numerous gulfs and bays
filled by lava plains at the scale of kilometers–tens of kilometers
(Fig. 18a,b). Both types of contacts indicate the stratigraphically
younger age of the plains.
Along the contacts of densely lineated plains with either shield or
regional plains, all lineaments of pdl stop at the contacts and none of
them penetrate into the surrounding plains (Fig. 18c,d). This abrupt
change of morphology across the contact indicates that densely
lineated plains have been partially flooded by materials of the vast
volcanic plains. The same features characterize the contacts of the
vast volcanic plains with ridged plains and ridge belts (Fig. 18e,f):
materials of the volcanic units deeply penetrate between the ridges
of unit pr/RB, which implies an older age for ridge belts.
The stratigraphic relationships of groove belts with the major
volcanic units are more complex. In the absolute majority of localities
where the belts and the plains units are in contact (Ivanov and Head,
2001a; Bridges and McGill, 2002; Stofan and Guest, 2003; McGill,
2004; Campbell and Clark, 2006), materials of either shield or
regional plains bury structures of the belts (Fig. 18g,h) indicating that
the plains are younger. In some other places, however, a few fractures/graben that continue to the general trend of the belts cut the
surface of the plains. These relationships suggest that deformation
within groove belts probably continued after emplacement of the
plains but was much less widespread and likely corresponded to the
waning phases of formation of the belts. Lobate plains always overlap
structures of groove belts.
Structures of rift zones cut the other tectonic units (t, pdl,
pr/RB, and gb) and the vast volcanic plains (psh, rp1, and rp2)
(Fig. 16a). Thus, rifts represent the stratigraphically youngest tectonic unit. Rifts often are in close spatial association with lobate
plains (Fig. 17). In areas where these units occur together, there is
strong evidence for both crosscutting of lobate plains by structures
of rifts and embayment of these structures by materials of the
plains (Fig. 16b). These relationships often occur in neighboring
areas and strongly suggest that rift zones and lobate plains formed
broadly simultaneously.
6. Discussion
6.1. Regimes of resurfacing
The definition of the tectonic and volcanic units (Fig. 1) and their
consistent global-scale relationships of relative ages provide the
possibility to outline the major features of the observable geologic
history of Venus. The class of tectonized terrains consists of tessera,
densely lineated plains, ridged plains/ridge belts, groove belts (all
these terrains comprise 20% of the surface of Venus, Table 2), and
rift zones ( 5% of the surface, Tables 1 and 2). Shield plains, regional
plains, and lobate plains represent the major volcanic units (Ivanov
and Head, 2013). Both shield plains and regional plains are the most
abundant ( 61% of the surface of Venus, Table 2) and ubiquitous
units on the planet (Ivanov and Head, 2011, 2013); lobate plains
make up 8% of the surface (Tables 1 and 2). The most widespread
tectonized terrains (t, pdl, pr/RB, and gb) are embayed by the vast
and mildly deformed volcanic units (Fig. 18) and, thus, are older. In
no place on Venus have we found the reverse stratigraphic relationships of the tectonized terrains with the vast plains (Ivanov and
Head, 2011, 2013). In contrast, structures of rift zones cut the surface
of regional plains in any regions where these units are in contact
(Fig. 16a). Rift zones are typically associated with lobate plains
(Fig. 17) and these units show relationships indicating their broadly
contemporaneous formation (Fig. 16b).
The globally observed stratigraphic relationships among the tectonic and volcanic units that make up the absolute majority ( 96%)
of the surface of Venus divide the observable portion of its geologic
history into three different episodes, each with a specific style of
resurfacing. These are as follows (Fig. 19): (1) Global tectonic regime,
when tectonic resurfacing dominated. Exposed occurrences of these
units comprise about 20% of the surface of Venus (Table 2). (2) Global
volcanic regime, when volcanism was the most important process of
resurfacing and resurfaced about 60% of Venus (Table 2). (3) Network
rifting-volcanism regime, when both tectonic and volcanic activities
were about equally important. During this regime, about 14% of the
surface of Venus was modified (Table 2).
Fig. 10. Spatial distribution of ridged plains/ridge belts plains (red) on Venus. Tessera massifs (light gray) and boundaries of the tessera clusters are shown for reference. The
background is the topographic map in simple cylindrical projection. The thicker black solid line corresponds to 0 km contour (6051 km), the thinner black solid line-contour
þ 1 km, and the thinner black dashed line-contour 1 km. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
21
impact craters. Lobate plains of the network-rifting and volcanism
regime embay 33% of craters, which suggests that during this
time the net volcanic intensity decreased and became comparable
to the rate of accumulation of craters.
6.2. The earlier global tectonic regime
Fig. 11. The age relationships of ridged plains and ridge belts (pr/RB) and tessera (t).
(a) Usually, materials of ridged plains embay pieces of tessera and structures of the
plains/belts cut tessera. Part of C1-MIDR 00N112, center of the image is at 1.9°N,
111.2°E. (b) In some areas (as, for example, at the NE flank of Ovda Regio), the usual
structures of ridge belts (white arrow) are cut by graben/fractures that extend from the
tessera massif (black arrows) and the structural pattern of the belt begins to resemble
that of tessera (tessera transitional terrain, ttt). Approximate extent of the ttt is shown
by dashed arrows. Part of C1-MIDR 00N095, center of the image is at 5.3°N, 98.8°E.
The mean crater density (craters ⩾8 km) on the surface of the vast
plains that define the global volcanic regime (psh, rp1, and rp2) is 1.83
per 106 km2 (Fig. 20), which is slightly larger than the mean crater
density for the whole planet, 1.69 per 106 km2. This means that the
tectonic terrains of the global tectonic regime (t, pdl, pr/RB, and gb),
which are embayed by the vast plains, occurred near the beginning
of the observable geological record of Venus (Fig. 19).
A statistical analysis based on the buffered crater density
(Kreslavsky et al., 2015) has shown that the tectonized terrains of
the global tectonic regime and the vast volcanic plains of the
global volcanic regime have similar crater retention ages. In contrast, units of the network-rifting and volcanism regime (lobate
plains and rift zones) have a significantly younger mean crater
retention age than that of regional plains. An analysis of truly
embayed craters on Venus (Ivanov and Head, 2015) reveals that
the proportion of craters embayed by regional plains is 3%. This
value is inconsistent with the predictions of the model of equilibrium resurfacing and suggests that during the global volcanic
regime volcanism acted in large regions and the process of formation of regional plains was more intensive than accumulation of
Relationships of relative ages between the ridges and graben in
tessera are crucially important for constraining of the possible
models of tessera formation. For example, in the upwelling or the
lava pond models (Hansen and Willis, 1996; Hansen et al., 1997;
Hansen, 2006) the graben (otherwise called ribbons (Pritchard
et al., 1997)) are considered as the oldest structures formed by
stretching and cracking of either the roof of the rising molten
diapir (Hansen and Willis, 1996) or “solidified scum of a huge lava
pond” (Hansen, 2006). The later thermal evolution of either the
diapir or the lava pond causes its cooling, subsidence, contraction
and deepening of the brittle–ductile transition, and formation of
the younger broad ridges (Hansen and Willis, 1996; Hansen, 2006).
In contrast, the downwelling models of tessera formation suggest
thickening of the crust over the sites of downwelling by the means
of under- and overthrusting, bending, and buckling of crustal slabs
that resulted in formation of broad contractional structures (specifically, ridges) at the beginning of tessera evolution (Bindschadler and Parmentier, 1990). The later viscous/brittle relaxation of
the crustal plateau may have caused formation of extensional
structures (graben) (Bindschadler and Parmentier, 1990; Bindschadler et al., 1992; Head et al., 1994; Gilmore et al., 1998).
The robust evidence for the older age of the tessera ridges
(Figs. 4–7) suggests that individual tesserae more likely represent
large and approximately equidimensional sites of contractional
structures formed primarily by compressional forces. This is completely consistent with the predictions of the downwelling model
and strongly contradicts the predictions of either the upwelling or
the lava pond hypotheses of tessera formation. Thus, it is likely that
tesserae formed in the tectonic environment of the large-scale contraction within several distinctive downwelling regions. The global
stratigraphic position of tessera suggests that this process apparently
was the most important at the very beginning of the visible (preserved) geologic history of Venus (Fig. 19).
Larger and smaller tessera massifs always occur as elevated
terrains (Fig. 3) and the largest tesserae form regional-scale highs
(Fig. 3, the right-side tail of the tessera hypsogram). This distinctive topographic signature of tessera terrain suggests that
(1) tessera occurrences are less susceptible to the embayment/
burial by later volcanic plains and (2) exposures of tessera mostly
mark regions of its formation. Thus, the tessera clusters, in which
tessera massifs preferentially occur (Fig. 8), may represent the loci
of downwelling. The dimensions of the clusters vary broadly from
about 2000 9000 km2 (Alpha–Lada cluster, A–L in Fig. 8) to
about 5000 14,000 km2 (Ovda–Thetis cluster, O–T in Fig. 8) and
may correspond to the upper limit of the sizes of the downwelling
cells. The major structural seams (e.g., long narrow troughs that
divide tesserae into series of blocks) characterize individual tessera regions but do not extend from one large tessera massif into
the other (Fig. 21). This suggests that individual large tesserae
probably evolved independently and their typical dimensions
(many hundreds to a few thousands of kilometers) may characterize the lower limit of the sizes of the downwelling cells.
In places where tessera is in contact with either densely lineated
plains (pdl) or/and ridged plains/ridge belts (pr/RB), the material
components of the later units usually embay tessera (Figs. 9a and 11a).
This implies their younger stratigraphic age. The density of impact
craters, however, is distinctly lower on tessera than that on the surface
of both pdl and pr/RB. About 15 craters larger than 16 km in diameter
are required to make the crater density on tessera to be at least equal
22
to that on the stratigraphically younger pdlþ pr/RB. This paradox of
the stratigraphic and crater ages suggests that craters were more
effectively erased from the surface of tessera probably due to continuing of tessera deformation. Two possible mechanisms may explain
this phenomenon. (1) The tectonic resurfacing can be more effective
within elevated regions due to possibly higher strain rates within the
areas that correspond to the thicker crust (Solomon, 1993). (2) A very
likely consequence of the downwelling model is formation of the deep
crustal roots beneath the growing tessera massifs. When the dragging
mantle forces waned, the roots of the lower density must begin to
float causing an epirogenetic uplift of the thickened crustal blocks.
Evidence for such uplift was found at the edges of Ovda Regio and
Alpha Tessera (Parker and Saunders, 1994; Head and Ivanov, 1996;
Gilmore and Head, 2000). Differential vertical and horizontal displacements of blocks during this stage of topographic evolution of the
larger tessera massifs could produce more deformation and partly
destroy the tessera population of impact craters. The amount of
epeirogenic uplift must be proportional to the depth of the crustal
roots. This suggests that topographically lower occurrences of ridged
plains/ridge belts (if they have crustal roots) could be displaced less
significantly compared with the higher tessera massifs, and their
crater record could be less affected.
Although materials of densely lineated plains and ridged plains
units usually embay tessera, in several regions relationships among
these units are more complex and are characterized by the tessera
transitional terrain, ttt (Figs. 9b and 11b). The most significant areas
of the tessera transitional terrain usually occur at the edges of the
large tesserae. In contrast to the tessera transitional terrain, pieces of
pdl and pr/RB that are away from tessera massifs, typically do not
show the intersections of the contractional and extensional structures and are deformed by one type of structure (Fig. 1). The preferential occurrence of the tessera transitional terrain near the larger
tessera massifs suggests continuous deformation within them and is
consistent with their late-stage epeirogenic uplift.
Ridged plains often form pronounced belts that can extend for
many hundreds of kilometers and consist of contractional structures
(e.g., Barsukov et al., 1986). This suggests that the belts formed under
compressional stresses applied within relatively narrow but very
extensive zones (Frank and Head, 1990; Kruychkov, 1992). These
characteristics of ridge belts resemble those of terrestrial thrust-and-
fold belts (e.g., van der Pluijm and Marshak, 2004), swarms of
wrinkle ridges on Mars (Golombek et al., 2001; Plescia, 1993; Montesi and Zuber, 2003; Mueller and Golombek, 2004), and lobate
scarps on Mercury (Watters et al., 2000), the structures of which are
interpreted to form over large thrust faults and indicate crustal
shortening due to lateral movements of crust/lithosphere. If similar
processes participated in formation of ridge belts on Venus (e.g.
Phillips et al., 1991), then the relatively low relief of the belts (a few
hundred meters (Young and Hansen, 2005)) suggests that the lateral
movements and related contraction were rather small (e.g., Solomon
et al., 1992), which characterizes displacements over the thrust faults
on Mars (Plescia, 1993).
The distinctive exceptions to this are the mountain belts
around Lakshmi Planum (Barsukov et al., 1986; Pronin, 1992). The
belts represent the highest mountain ranges on Venus and their
relationships with the surrounding terrains provide evidence for
large-scale shortening, collision, underthrusting, and epeirogenic
uplift (Head, 1990; Vorder Bruegge and Head, 1991; Ivanov and
Fig. 13. Branches of groove belts (gb) usually form rims of coronae. Because of this,
coronae and groove belts are often spatially associated with each other. Part of C1MIDR 15S215, center of the image is at 20.8°S, 220.2°E.
Fig. 12. Spatial distribution of groove belts plains (red) on Venus. Tessera massifs (light gray) and boundaries of the tessera clusters are shown for reference. The background
is the topographic map in simple cylindrical projection. The thicker black solid line corresponds to 0 km contour (6051 km), the thinner black solid line-contour þ1 km, and
the thinner black dashed line-contour 1 km. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
23
Fig. 14. Spatial distribution of groove belts (black areas) and coronae. (a) Coronae of type 1 are shown as red circles; inset shows the bimodal longitudinal distribution of
coronae. Coronae of type 2 are shown as green stars. Both types of coronae are in close spatial association with groove belts. (b) Map of the spatial distribution of groove belts
centered at 0°E shows the presence of two huge concentrations of the belts. All maps are in simple cylindrical projection. (For interpretation of the references to color in this
figure legend, the reader is referred to the web version of this article.)
Head, 2008b). The mountain belts, however, exist only in one
region, which suggests that even if the belts are related to processes akin to subduction, they were fairly restricted on Venus.
Pieces of ridged plains tend to occur within and at the periphery
of the major tessera clusters and prominent zones of ridge belts are
often aligned conformably with large tessera massifs (Fig. 10). Near
the northern and northwestern edges of Ovda Regio ridge belts are
merging with tessera. In this area, structures of the belts are parallel
to the tessera ridges and in places are cut by tessera graben (Fig. 11b).
Thus, the belts seem to build on tessera terrain in northern Ovda. In
several other regions on Venus, ridge belts also appear to be incorporated into tessera (e.g., Semuni Dorsa at NW flank of Fortuna
Tessera (Ivanov and Head, 2007), Allat Dorsa at the northern edge of
Dekla Tesserae (Ivanov and Head, 2008b)). The same nature (contractional ridges) and parallelism of the major structures in the belts
and tesserae, but obviously lower density of tectonic structures
within the belts, suggest that the belts may represent edge facies of
the large downwelling sites and correspond to the waning phases of
downwelling.
Thus, the earlier phases of the global tectonic regime were
characterized by the predominance of contractional structures
(tessera ridges, ridge belts, mountain belts) that were likely to
have been related to lateral movements of the lithosphere.
In comparison with the other terrains of the global tectonic
regime, groove belts are more broadly distributed over the surface
and do not show evidence of association with the tessera clusters
(Fig. 12). The consistent relative age relationships of groove belts
imply that the belts formed during the later phases of the global
tectonic regime. The nature of structures of the belts implies that
these phases were predominantly related to broadly distributed
extension of the crust/lithosphere (Fig. 19).
An important characteristic of groove belts is that their branches
often represent the tectonic components of coronae (Fig. 13) that
form chains interconnected by segments of the belts (Ivanov and
24
Fig. 15. Structures of groove belts (gb) cut the surface of tessera (t, black arrows)
and ridged plains/ridge belts (pr/RB, white arrows) and, thus, are younger. These
types of stratigraphic relationships are seen in all regions where these units occur
together. Part of C1-MIDR 45N096, center of the image is at 50.8°N, 86.9°E.
Head, 2004a, 2012). Coronae are thought to be the surface manifestations of mantle diapirs (e.g., Pronin and Stofan, 1990; Stofan and
Head, 1990; Stofan et al., 1991, 1992; Smrekar and Parmentier, 1996;
Smrekar and Stofan, 1997, 1999; Musser and Squyres, 1997; Herrick,
1999; Jellinek et al., 2002). Thus, the close spatial association of the
belts and coronae (Fig. 14a,b) suggest that these features formed
mutually due to multiple and relatively small-scale mantle upwellings/diapirs (84% of all coronae are in a diameter range of 100–
400 km). The possible nature of coronae as the surface manifestation
of mantle diapirs suggests that coronae are likely to have served as
important magmatic centers, manifested by lava flows commonly
emanating from coronae and radiating swarms of graben that are
often associated with coronae (Grosfils and Head, 1994, 1995, 1996;
Studd et al., 2010, 2011). The graben of the swarms are interpreted as
the surface manifestation of dikes sourced by the magmatic reservoirs beneath some coronae (e.g., Grosfils and Head, 1994; Ernst
et al., 1995, 2001, 2003). In places, the radiating graben merge with
groove belts that are associated with coronae. Thus, the belts may
contain a volcano-tectonic component related to evolution of the
coronae magmatic centers.
At the global scale, the branches and segments of groove belts
appear to be organized into two giant circular clusters (about
15,000 km in diameter each) that are centered at the equator and
at about 75°E and 285°E respectively (Fig. 14b). The branches of
groove belts are topographically elevated features (Fig. 3) and thus
are less susceptible to later flooding by the younger volcanic
plains. This suggests that the contiguous areas in the northern and
southern hemispheres and in a broad longitudinal zone at 180°E
(Fig. 14b) where the belts are absent are not due to the later volcanic embayment/burial but indicate true gaps in the spatial distribution of groove belts. The well-known bimodal longitudinal
distribution of coronae (Stofan et al., 1992) (Fig. 14a, inset) probably also reflects the existence of the clusters of groove belts.
Different types of large-scale features are spatially associated
with the groove belt clusters. The majority of the largest tessera
massifs occur in the western cluster, and the Beta–Atla–Themis
(BAT) region is at the edge of the eastern one (Fig. 14b). Within the
western cluster, coronae of type 1 (Stofan et al., 2001) are more
evenly distributed throughout its territory, whereas in the eastern
cluster they are strongly concentrated within the BAT region
(Fig. 14a). Coronae of type 2 that are characterized by an incomplete rim (less than 180° arc) and strongly shifted toward the
Fig. 16. Morphology and age relationships of rift zones (rz). This tectonic unit
represents swarms of densely packed, straight to curvilinear graben and canyons
that can reach tens of kilometers in width and hundreds of kilometers in length.
(a) When rift zones (rz) cross expances of regional plains (rp1), structures of rifts
cut the surface of the plains and, thus, postdate this unit and all older units. Part of
C1-MIDR 15N197, center of the image is at 12.2°N, 199.3°E. (b) Rift zones and lobate
plains (pl) show relationships that suggest their broadly contemporaneous formation: in some places structures of rifts cut the plains (black arrows) and in a
neighboring region lobate plains embay structures of rift zones (white arrows). Part
of C1-MIDR 00N197, center of the image is at 6.2°N, 199.8°E.
“rimmed depression” and “rim only” topographic groups (Stofan
et al., 2001) are thought to correspond to the late stages of evolution of the corona-related mantle diapirs (Smrekar and Stofan,
1997). Coronae of type 2 are spatially associated with segments of
groove belts but do not show a tendency to preferentially occur
within any of the groove belt clusters (Fig. 14a).
The largest corona-like feature on Venus, Artemis Corona, is at
the edge of the western cluster of groove belts. Although Artemis
is about 2500 km in diameter and could potentially affect the
spatial distribution of groove belts in its broad surroundings, the
general pattern of the belts does not indicate the presence of
Artemis (Fig. 14). In contrast, the general pattern of wrinkle ridges
within broad plains southward of Artemis is aligned conformably
with both the southern edge of eastern Aphrodite and Artemis
(Bilotti and Suppe, 1995, 1999; Hansen and Olive, 2010). Such an
alignment was interpreted either as the result of the swell-push
body force that is applied to the lithosphere by elevated region of
the eastern Aphrodite or as the result of evolution of Artemis due
to “coupling of convective mantle flow with the lithosphere”
(Hansen and Olive, 2010). In any case, the major episodes of
25
Fig. 17. Spatial distribution of rift zones (red) and lobate plains (green) on Venus. Tessera massifs (light gray) and boundaries of the tessera clusters are shown for reference.
The background is the topographic map in simple cylindrical projection. The thicker black solid line corresponds to 0 km contour (6051 km), the thinner black solid linecontour þ1 km, and the thinner black dashed line-contour 1 km. (For interpretation of the references to color in this figure legend, the reader is referred to the web
version of this article.)
Fig. 18. a,b Age relationships between the major tectonic and volcanic units. Materials of shield plains (psh) and regional plains (rp) embay tessera (t). (a) Part of C1-MIDR
15S060, center of the image is at 9.3°S, 54.5°E. (b) Part of C1-MIDR 15S060, center of the image is at 9.3°S, 54.5°E. (c, d) Age relationships between the major tectonic and
volcanic units. Materials of shield plains (psh) and regional plains (rp) embay densely lineated plains (pdl). (a) Part of C1-MIDR 45S265, center of the image is at 39.4°S,
262.8°E. (b) Part of C1-MIDR 45N307, center of the image is at 42.4°N, 310.7°E. (e, f) Age relationships between the major tectonic and volcanic units. Materials of shield
plains (psh) and regional plains (rp) embay ridged plains/ridge belts (pr/RB). (a) Part of C1-MIDR 45N159, center of the image is at 37.8°N, 156.6°E. (b) Part of C1-MIDR
15N060, center of the image is at 9.4°N, 65.9°E. (g, h) Age relationships between the major tectonic and volcanic units. Materials of shield plains (psh) and regional plains (rp)
embay groove belts (pr/RB). (a) Part of C1-MIDR 30N243, center of the image is at 34.2°N, 249.5°E. (b) Part of C1-MIDR 45S350, center of the image is at 40.8°S, 344.2°E.
formation of Artemis probably postdated groove belts and predated formation of wrinkle ridges.
Groove belts are younger than the tectonic terrains that may be
indicative of limited lateral displacement of lithosphere (e.g., ridge
belts). In no region on Venus, however, either isolated contractional
structures or their swarms, which are complementary to groove belts
both stratigraphically and by the scale, were observed. The lack of
such structures/zones suggests diminishing of lateral movements of
lithosphere and re-orientation of tectonic deformation mostly to
vertical movements. The tight spatial and stratigraphic association of
groove belts with coronae may indicate that the dominant style of
resurfacing of the later phases of the global tectonic regime was
plume-tectonics that caused the predominantly vertical displacement of lithosphere and its deformation.
26
6.3. The global volcanic regime
The earlier tectonic regime was followed by emplacement of
the vast volcanic plains (shield and regional plains), the surface of
which is pervasively deformed by the global network of wrinkle
ridges (Bilotti and Suppe, 1999). Emplacement of the plains has
defined the second, globally volcanic regime, when tectonic
deformation related to mantle convection waned and volcanic
resurfacing dominated (Fig. 19).
Details of the global volcanic regime have been described in a
separate paper (Ivanov and Head, 2013). Here we emphasize that
the density of impact craters on terrains of the earlier tectonically
dominated regime (t, pdl, pr/RB, and gb) and on vast volcanic
plains (psh and rp) are practically indistinguishable (Fig. 20). The
stratigraphic relationships between these terrains, however, are
always clear and unambiguously indicate the younger age of the
plains. This suggests a rapid change from the earlier global tectonic
regime to the following global volcanic resurfacing.
Table 2
Areas modified during different regimes of resurfacing.
Unit
Unit area, 106 km2
Unit area, %
Global tectonic regime
t
pdl
pr
gb
Sum
35.7
7.8
10.3
39.9
93.6
7.8
1.7
2.2
8.7
20.3
Global volcanic regime
psh
rp1
rp2
Sum
85.2
152.0
45.2
282.4
18.5
33.0
9.8
61.4
Network rifting-volcanism regime
rz
24.3
pl
40.7
Sum
65.0
5.3
8.8
14.1
Fig. 20. Mean crater densities on units that formed during different regimes of
resurfacing on Venus. The densities for the global tectonic and the global volcanic
regimes are statistically indistinguishable, but the crater density on rift zones and
lobate plains (the network rifting-volcanic regime) is much smaller. Note that error
bars of the density estimates are 4 sigmas.
Fig. 19. A global correlation chart that shows the three major regimes of resurfacing on Venus: global tectonic regime, global volcanic regime, and Network-rifting-volcanism
regime (see text for discussion).
27
Fig. 21. Spatial distribution of the most prominent tectonic seams (narrow and long troughs) within tessera massifs. The pattern of the seams suggests that the tessera
massifs represented discontinuous regions.
Within areas of both regional and shield plains there is little
evidence for either impact craters that are obviously embayed (e.g.,
crater Heloise, Fig. 22a) or structures resembling ghost craters
(Fig. 22b). One may argue that coronae of type 2 represent impact
structures completely covered by lava plains (the ghost craters). The
mean diameter of these coronae is 240 km (7130 km) (Stofan
et al., 2001). This suggests that if coronae of type 2 are indeed flooded craters then only larger impact structures contributed to their
population, which is not very realistic. Another argument is that the
smaller craters could be completely hidden by lava flows/plains
without any evidence of their presence remaining. Ghost craters
caused by flooding and burial of impact craters by lava are common
features on Mars, Mercury and the Moon. On these planets, the ghost
craters are seen in volcanically resurfaced regions where the total
thickness of overlying volcanic layers is estimated to be 0.5–1.5 km
(Ivanov et al., 2005; Head et al., 2011; Klimczak et al., 2012). Thus, a
few kilometers of lavas may be needed to effectively hide evidence
for the presence of impact craters several tens of kilometers in
diameter.
Several lines of evidence suggest, however, that the mean
thickness of plains that have been formed during the global volcanic regime is not that large.
(1) In some places, ghost tectonic structures are seen within areas
covered by shield/regional plains, which also indicates a small
thickness of the plains, a few hundred meters (DeShon et al.,
2000).
(2) Occurrences of densely lineated plains are broadly distributed
within regional plains (Fig. 8). Clusters of the pdl pieces are seen
near the unflooded highlands as well as within large extensions
of regional/shield plains away from the highlands (e.g., in Atalanta or Sedna Planitiae, Fig. 8). Because relief of occurrences of
pdl is small (Ivanov and Head, 2001a) and their surface appears
to be gently undulating without abrupt changes in elevation, the
existence of clusters of pieces of pdl requires a rather small
thickness of embaying/flooding materials. A layer of greater
thickness, for example, a few kilometers would effectively hide
exposures of densely lineated plains.
(3) The average slope of the surface of shield and regional plains
and the underlying units are small and kipukas of underlying
units are seen in broad zones along the contact between the
plains units and older tectonized terrains. Collins et al. (1999)
have estimated the thickness of the overlying regional plains
in the transitional zone to be about 0.5 km. The broad spatial
distribution of the occurrences of t, pdl, pr and gb (Figs. 2, 8,
10, and 12) indicates that these transitional zones are globally
extensive and characterize more than a half of the surface of
Venus.
(4) Near the contact with shield plains, kipukas of small volcanoes
are seen within regional plains suggesting the thickness of the
plains in these areas to be less than 100–200 m (Kreslavsky
and Head, 1999).
Thus, the units of the global volcanic regime do not appear to be
thick enough (at least in broad transition zones from them to the
older tectonized terrains) to hide completely the preceding crater
record. This suggests that the volcanic units of shield and regional
plains were likely to have been emplaced onto a surface that was
largely free of craters and, thus, the previous craters have been
mostly erased during the periods prior to the global volcanic regime.
6.4. The network rifting-volcanism regime
The network rifting-volcanism regime characterized the later
episodes of the geologic history of Venus; rift zones represent a
tectonic component of this regime and lobate plains represent a
volcanic component (Fig. 19).
Rift zones are very prominent morphologically but their total area
is about four times smaller than the exposed (i.e., minimal) area of
tectonic terrains of the global tectonic regime (Table 1). In contrast to
these earlier terrains, which are broadly distributed over the surface
of Venus, rift zones form a few very long and pronounced zones that
are prominently concentrated in the equatorial and the BAT regions
of Venus (Figs. 1 and 23a). These characteristics of rifts suggest that
(1) the effect of tectonic resurfacing diminished with time throughout the visible geologic history and (2) the style of tectonic resurfacing evolved from the broadly distributed deformation near the
beginning of the visible geologic history to the highly concentrated
deformation during the late episodes of history. These changes in
both the intensity and the lateral extent of tectonic resurfacing are
consistent with the transition from an earlier mobile-lid regime of
mantle circulation with thinner/weaker lithosphere (e.g., Moresi and
28
Fig. 22. (a) Example of a crater (arrow) that is heavily embayed by materials of
regional plains (both the lower and upper units, rp). Crater Heloise, center of the
image is at 40.1°N, 52.0°E. (b) Example of a possible ghost crater (arrows). This
feature is represented by a low circular topographic rim about 20 km in diameter.
The SW segment of the rim is absent and this portion of the structure is cut by a
lava channel (portion of Baltis Vallis). Center of the image is at 48.7°N, 163.0°E.
Solomatov, 1998; Phillips and Hansen, 1998; Brown and Grimm,
1999) to a later stagnant-lid regime with thicker/stronger lithosphere
(e.g., Parmentier and Hess, 1992; Head et al., 1994; Solomatov and
Moresi, 1996; Grimm and Hess, 1997).
The areal distribution of rift zones reveals a significant reorganization of the pattern of major zones of extension during the
network rifting-volcanism regime (Fig. 23). The segments of the
older groove belts, although they are concentrated near Beta, Atla,
and Themis Regiones, do not interconnect them (Fig. 14b). This
may mean that the BAT region began its evolution during the late
phases of the global tectonic regime (formation of groove belts) as
a series of individual hotspots. The high topography, general
dome-like shape, gravity signatures, and abundant young volcanism of Beta, Atla, and Themis Regiones collectively suggests that
these rises are related to the large-scale and likely deep mantle
upwelling/plumes (e.g., Esposito et al., 1982; Smrekar and Phillips,
1991; Simons et al., 1994; Smrekar et al., 1997). In contrast to
groove belts, the branches of rift zones interconnect Beta, Atla, and
Themis and clearly outline the BAT region (Fig. 23a) suggesting
that during the network rifting-volcanism regime the individual
mantle plumes of the BAT region began to interact with each other
and control the spatial distribution of regional extensional
deformation (rift zones). A thicker and maybe stronger thermal
lithosphere could provide a necessary medium for the interaction
between the remote centers of uplift during this time. Measurements of the spacing of structures within groove belts and rift
zones have shown (Mastrapa, 1997; Guseva, 2008) that the spacing
is greater in rifts reflecting the greater thickness of the deformed
layer (Ladeira and Price, 1981; Wu and Pollard, 1995; Ji and Saruwatari, 1998).
Another possible line of evidence for the thicker thermal lithosphere during the network rifting-volcanism regime is that rift zones
are poorly correlated with coronae but are clearly associated with the
large dome-shaped rises of the BAT region (Fig. 23b). This correlation
suggests that during this time the large-scale mantle upwellings/
plumes (likely manifested by the rises) controlled the spatial distribution of the principal zones of extension and rupture of the crust.
The smaller-scale mantle diapirs (likely manifested by coronae) have
played a subordinate role in resurfacing during the network riftingvolcanism regime. This can be another consequence of the thicker/
stronger lithosphere that may have acted as a rheological barrier that
filtered out smaller mantle diapirs but was more “transparent” to the
larger mantle upwellings/plumes.
In summary, the network rifting-volcanism regime (Atlian
period, Fig. 19) was characterized by both a significant diminishing
of tectonic and volcanic activity (Table 2) and a concentration of
tectonism and volcanism within fewer distinctive regions/zones.
Fig. 20 shows that the mean density of craters superposed on
the surface of both rift zones and lobate plains is significantly
lower than the crater densities on units of the global tectonic and
volcanic regimes (Ivanov and Head, 2015; Kreslavsky et al., 2015).
This suggests that the transition from the global volcanic regime to
the fully developed network rifting-volcanism regime was longer
than the change from the global tectonic to global volcanic
regimes whose units are indistinguishable on the basis of crater
density (Fig. 20). The apparently abrupt change from the tectonically to volcanically dominated styles of resurfacing may correspond to a rapid change in mantle convection, whereas the
extended transition from the global volcanic to the network rifting-volcanism regimes may reflect a gradual increase of the
lithosphere thickness.
7. Summary/conclusions
The scale and abundance of tectonic features related to the postemplacement deformation of materials provide the possibility to
define two major types of endogenous terrains of Venus: volcanic
and tectonic units. Tectonic structures dominate the surface of the
tectonic units and erase the original morphologic characteristics of
the underlying materials. In this respect, tectonic units are distinctly
different from individual structures that, although are abundant in
places, do not obscure the initial nature of volcanic materials (e.g.,
wrinkle ridges, radiating swarm of graben, individual fractures, etc.).
The tectonic units defined in such a way are (1) tessera (t), (2) densely lineated plains (pdl), (3) ridged plains/ridge belts (pr/RB),
(4) groove belts (gb), and (5) rift zones (rz). These units comprise
about 25.5% of the surface of Venus.
Structures that characterize the tectonic units reflect different
styles of tectonism. Tessera and ridged plains are related to compressional forces whereas densely lineated plains, groove belts,
and rift zones manifest the dominance of tensional environments.
The stratigraphic relationships among the tectonic units suggest
that tessera, densely lineated plains, and ridged plains were
formed near the very beginning of the visible portion of the geologic history of Venus. Tessera appears as the oldest terrain but pdl
and pr/RB were overlapping with it in time, which is suggested by
the tessera transitional terrain that bears structural characteristics
29
Fig. 23. (a) Spatial distribution of rift zones clearly outline the Beta–Atla–Themis (BAT) region. The pattern of the distribution of the older tectonic units did not show the
presence of BAT (see. Fig. 14c). (b) Neither coronae of type 1 (red circles), nor coronae of type 2 (green stars) show spatial associations with rift zones. (For interpretation of
the references to color in this figure legend, the reader is referred to the web version of this article.)
of tessera, pdl, or pr/RB. Groove belts postdate these tectonized
terrains but are older than the vast volcanic plains such as shield
plains and regional plains. Rift zones are among the youngest
terrains on Venus.
The stable and globally consistent relationships of relative ages
between the tectonic and volcanic units allow division of the
observable portion of the geologic history into three parts, each
with its specific style of resurfacing. The oldest part, the global
tectonic regime, is defined by such units as t, pdl, pr/RB and gb.
This time span is apparently divided into the earlier phase when
compressional forces responsible for the formation of tessera and
ridged plains dominated and the later phase when globally distributed extension formed groove belts. During the middle part,
the global volcanic regime, vast volcanic plains (shield and regional plains) were emplaced. Tectonic structures mildly deform the
surface of these units and volcanism was the principal contributor
to the resurfacing during this time (Ivanov and Head, 2013).
Approximately equally abundant volcanic and tectonic processes
that have resurfaced about 15% of the surface of Venus characterize the youngest part of the geologic history, the network
rifting-volcanism regime.
The crater densities on the surfaces formed and modified during the global tectonic and global volcanic regimes are indistinguishable. However, the stratigraphic relationships between the
units of the global tectonic regime and plains of the global volcanic
regime always indicate the younger relative age of the plains. This
suggests a rapid change from the earlier global tectonic regime to
the following global volcanic resurfacing. In contrast, the mean
density of craters superposed on the surfaces formed during the
global network rifting-volcanism regime is significantly lower
than the crater densities on units of the global volcanic regime.
This suggests that the transition to the fully developed network
rifting-volcanism regime was longer than the transition between
both the global tectonic and volcanic regimes.
30
Ridges of tesserae and ridge belts suggest that during the earlier phases of the global tectonic regime compressional forces
dominated. They were responsible for the formation of bulk of
tesserae over the sites of mantle downwelling and ridge belts due
to limited horizontal displacement and warping/buckling of crustal materials at the periphery of the major downwelling cells.
Groove belts characterize the later phase of the global tectonic
regime and are spatially associated with coronae. The contractional
structures that are complementary to groove belts both stratigraphically and by scale are not observed. This suggests that during
the later phase of the global tectonic regime vertical displacement
and rupture of lithosphere prevailed and was caused by a plumedominated tectonic style. Chains of coronae interconnected by
groove belts are the manifestations of this type of tectonism.
The major tectonic structures of the global volcanic regime are
wrinkle ridges and graben swarms radiating from distinct volcanotectonic centers. Both types of structures deform the surfaces of the
vast plains but do not conceal the original morphologies of the plains
and are not related directly to specific patterns of mantle circulation.
Rift zones that constitute the tectonic component of the network rifting-volcanism regime are very prominent morphologically but their total area is about four times smaller than the
exposed area of tectonic terrains of the global tectonic regime. This
means that (1) the effect of tectonic resurfacing diminished with
time throughout the visible geologic history of Venus and (2) the
style of tectonic resurfacing evolved from the broadly distributed
deformation during the earlier global tectonic regime to the highly
concentrated deformation during the later network rifting-volcanism regime. The nature of rift zones and their tight spatial
association with the large dome-shaped rises suggest that formation of rifts and, thus, the tectonic style during the network
rifting-volcanism regime was due to localized zonal lithospheric
disruption over the sites of large mantle upwellings that may
resemble the environments of the terrestrial continental rifts
(McGill et al., 1981; Stofan et al., 1989).
Thus, these broad outlines of the geologic history of tectonism,
together with the geological history of volcanism (Ivanov and Head,
2013) and the global geological map of Venus (Ivanov and Head,
2011), provide the basis to help distinguish among the range of
proposed models for the geodynamic evolution of Venus (Shalygin
et al., 2012; McGovern et al., 2013; Driscoll and Bercovici, 2014;
Gillmann and Tackley, 2014; Taylor, 2014; Campbell et al., 2015).
Acknowledgments
We gratefully acknowledge fruitful and productive discussions
with Alexander Basilevsky and numerous participants in the
Venus quadrangle mapping effort sponsored by NASA and managed by the U.S. Geological Survey. We gratefully acknowledge the
detailed reviews from two anonymous reviewers.
References
Addington, E.A., 2001. A stratigraphic study of small volcano clusters on Venus.
Icarus 149, 16–36.
Arkani-Hamed, J., Toksoz, M.N., 1984. Thermal evolution of Venus. Phys. Earth
Planet. Inter. 34, 232–250.
Arvidson, R.E., Baker, V.R., Elachi, C., Saunders, R.S., Wood, L.A., 1991. Magellan:
initial analysis of Venus surface modification. Science 252, 270–275.
Arvidson, R.E., Greeley, R., Malin, M.C., et al., 1992. Surface modification of Venus as
inferred from Magellan observation of plains. J. Geophys. Res. 97, 13303–13317.
Aubele, J.C., Slyuta, E.N., 1990. Small domes on Venus: characteristics and origin.
Earth Moon Planets 50/51, 493–532.
Avduevsky, V.S., Marov, M.Ya., Kulikov, Yu.N., Shari, V.P., Gorbachevsky, A.Ya.,
Uspensky, G.R., Cheremukhina, Z.P., 1983. Structure and parameters of the
Venus atmosphere according to Venera probe data In: Hunten, D.M., Colin, L.,
Donahue, V.I., Moroz, V.I. (Eds.), Venus. University Arizona Press, Tucson,
pp. 280–298.
Barsukov, V.L., 1992. Venusian igneous rocks In: Barsukov, V.L., Basilevsky, A.T.,
Volkov, V.P., Zharkov, V.N. (Eds.), Venus Geology, Geochemistry, and Geophysics
(Research Results from the USSR). University of Arizona Press, Tucson, London,
pp. 165–176.
Barsukov, V.L., Basilevsky, A.T., Burba, G.A., et al., 1986. The geology and geomorphology of the Venus surface as revealed by the radar images obtained by
Venera 15 and 16. J. Geophys. Res. 91, D399–D411.
Basilevsky, A.T., 2008. Geologic Map of the Beta Regio Quadrangle (V–17), Venus.
USGS Scientific Investigations Map 3023
Basilevsky, A.T., Head, J.W., 1995a. Global stratigraphy of Venus: analysis of a random sample of thirty-six test areas. Earth Moon Planets 66, 285–336.
Basilevsky, A.T., Head, J.W., 1995b. Regional and global stratigraphy of Venus: a
preliminary assessment and implications for the geological history of Venus.
Planet. Space Sci. 43, 1523–1553.
Basilevsky, A.T., Head, J.W., 1998. The geologic history of Venus: a stratigraphic
view. J. Geophys. Res. 103, 8531–8544.
Basilevsky, A.T., Head, J.W., 2000. Rifts and large volcanoes on Venus: global
assessment of their age relations with regional plains. J. Geophys. Res. 105,
24583–24661.
Bilotti, F., Suppe, J., 1995. A global analysis of the distribution of wrinkle ridges on
Venus. EOS Trans. AGU 74, F333.
Bilotti, F., Suppe, J., 1999. The global distribution of wrinkle ridges on Venus. Icarus
139, 137–157.
Bindschadler, D.L., Parmentier, E.M., 1990. Mantle flow tectonics: the influence of a
ductile lower crust and implications for the formation of topographic uplands
on Venus. J. Geophys. Res. 95, 21329–21344.
Bindschadler, D.L., Head, J.W., 1991. Tessera terrain, Venus: characterization and
models for origin and evolution. J. Geophys. Res. 96, 5889–5907.
Bindschadler, D.L., Schubert, G., Kaula, W.M., 1992. Coldspots and hotspots: global
tectonics and mantle dynamics of Venus. J. Geophys. Res. 97, 13495–13532.
Bridges, N.T., 2002. G.E. McGill Geologic Map of the Kaiwan Fluctus Quadrangle (V44), Venus. USGS Scientific Investigations Map 2747
Brown, C.D., Grimm, R.E., 1999. Recent tectonic and lithospheric thermal evolution
of Venus. Icarus 139, 40–48.
Campbell, B.A., Campbell, P.G., 2002. Geologic Map of the Bell Regio Quadrangle (V9), Venus. USGS Scientific Investigations Map 2743
Campbell, B.A., Clark, D.A., 2006. Geologic Map of the Mead Quadrangle (V-21),
Venus. USGS Scientific Investigations Map 2897
Campbell, B.A., Campbell, D.B., Morgan, G.A., Carter, L.M., Nolan, M.C., Chandler, J.F.,
2015. Evidence for crater ejecta on Venus tessera terrain from Earth-based
radar images. Icarus 250, 123–130.
Collins, G.C., Head, J.W., Basilevsky, A.T., Ivanov, M.A., 1999. Evidence for rapid
regional plains emplacement on Venus from the population of volcanically
embayed impact craters. J. Geophys. Res. 104, 24121–24140.
Collins, G.C., Patterson, G.W., Head, J.W., Pappalardo, R.T., Prockter, L.M., Lucchitta, B.K., Kay, J.P., 2013. Global Geologic Map of Ganymede. USGS Scientific
Investigations Map 3237
Copp, D.L., Guest, J.E., 2007. Geologic Map of the Sif Mons Quadrangle (V-31),
Crumpler, L.S., Aubele, J., 2000. Volcanism on Venus In: Houghton, B., Rymer, H.,
Stix, J. (Eds.), Encyclopedia of Volcanoes. Academic Press, San Diego, San
Francisco, New York, Boston, London, Sydney, Toronto, pp. 727–770.
DeShon, H.R., Young, D.A., Hansen, V.L., 2000. Geologic evolution of southern
Rusalka Planitia. J. Geophys. Res. 105, 6983–6996.
Driscoll, P., Bercovici, D., 2014. On the thermal and magnetic histories of Earth and
Venus: influences of melting, radioactivity, and conductivity. Phys. Earth Planet.
Inter. 236, 36–51.
Ernst, R.E., Head, J.W., Parfitt, E., Grosfils, E., Wilson, L., 1995. Giant radiating dyke
swarms on Earth and Venus. Earth Sci. Rev. 39, 1–58.
Ernst, R.E., Grosfils, E.B., Mege, D., 2001. Giant dike swarms: Earth, Venus, and Mars.
Annu. Rev. Earth Planet. Sci. 29, 489–534.
Ernst, R.E., Desnoyers, D.W., Head, J.W., Grosfils, E.B., 2003. Graben-fissure systems
in Guinevere Planitia and Beta Regio (264–312E, 24–60N), Venus, and implications for regional stratigraphy and mantle plumes. Icarus 164, 282–316.
Esposito, P.B., Sjogren, W.L., Mottinger, N.A., Bills, B.G., Abbot, E., 1982. Venus
gravity: analysis of Beta Regio. Icarus 51, 448–459.
Fernandez, C., Anguita, F., Ruiz, J., Romeo, I., Martin-Herrero, A.I., Rodriguez, A.,
Pimentel, C., 2010. Structural evolution of Lavinia Planitia, Venus: implications
for the tectonics of the lowland plains. Icarus 206, 210–228.
Frank, S.L., Head, J.W., 1990. Ridge belts on Venus: morphology and origin. Earth
Moon Planets 50/51, 421–470.
Ghent, R., Hansen, V.L., 1999. Structural and kinematic analysis of Eastern Ovda
Regio, Venus: implications for crustal plateau formation. Icarus 139, 116–136.
Gillmann, C., Tackley, P., 2014. Atmosphere/mantle coupling and feedbacks on
Venus. J. Geophys. Res. 119, 1189–1217. http://dx.doi.org/10.1002/
2013JE004505.
Gilmore, M.S., Head, J.W., 2000. Sequential deformation of plains at the margins of
Alpha Regio, Venus: implications for tessera formation. Meteorit. Planet. Sci. 35,
667–687.
Gilmore, M.S., Ivanov, M.A., Head, J.W., Basilevsky, A.T., 1997. Duration of tessera
deformation on Venus. J. Geophys. Res. 102, 13357–13368.
Gilmore, M.S., Collins, G.C., Ivanov, M.A., Marinangeli, L., Head, J.W., 1998. Style and
sequence of extensional structures in tessera terrain, Venus. J. Geophys. Res.
103, 16813–16840.
Gilmore, M.S., Resor, P.G., Ghent, R., Senske, D.A., Herrick, R.R., 2010. Mapping and
modeling of a tessera collision zone, Tellus Regio, Venus. In: Proceedings of the
41st Lunar and Planetary Science Conference, Abstract 1769.
Golombek, M.P., Anderson, F.S., Zuber, M.T., 2001. Martian wrinkle ridge topography: evidence for subsurface faults from MOLA. J. Geophys. Res. 106,
23811–23821.
Graupner Bergmann, M., Hansen, V.L., 2013. Structural and geologic mapping of
southern Tellus Regio, Venus. In: Proceedings of the 44th Lunar and Planetary
Science Conference, Abstract 1542.
Greeley, R., Guest, J.E., 1987. Geologic Map of the Eastern Equatorial Region of Mars,
1:15,000,000 Scale USGS Scientific Investigations Map 1802-B
Grimm, R.E., Hess, P.C., 1997. The crust of Venus In: Bougher, S.W., Hunten, D.M.,
Phillips, R.J. (Eds.), Venus II Geology, Geophysics, Atmosphere, and Solar Wind
Environment. University Arizona Press, Tucson, pp. 1163–1204.
Grosfils, E.B., Head, J.W., 1994. The global distribution of giant radiating dike
swarms on Venus: implications for the global stress stare. Geophys. Res. Lett.
21, 701–704.
Grosfils, E.B., Head, J.W., 1995. Radiating dike swarms on Venus: evidence for
emplacement at zones of neutral buoyancy. Planet. Space Sci. 43, 1555–1560.
Grosfils, E.B., Head, J.W., 1996. The timing of giant radiating dike swarm emplacement on Venus: implications for resurfacing of the planet and its subsequent
evolution. J. Geophys. Res. 101, 4645–4656.
Grosfils, E.B., Long, S.M., Venechuk, E.M., Hurwitz, D.M., Richards, J.W., Kastl, B.,
Drury, D.E., Hardin, J., 2011. Geologic Map of the Ganiki Planitia Quadrangle (V14), Venus USGS Scientific Investigations Map 3121.
Guest, J.E., Bulmer, M.H., Aubele, J., Beratan, K., Greeley, R., Head, J.W., Michaels, G.,
Weitz, C., Wiles, C., 1992. Small volcanic edifices and volcanism in the plains of
Venus. J. Geophys. Res. 97, 15949–15966.
Guseva, E.N., 2008. Comparative analysis of topography of the Venusian rifts and
terrestrial continental rifts in Africa. In: Proceedings of the 39th Lunar and
Planetary Science Conference, Abstract 1063.
Hansen, V.L., 2006. Geologic constraints on crustal plateau surface histories, Venus:
the lava pond and bolide impact hypotheses. J. Geophys. Res. 111, E11010. http:
//dx.doi.org/10.1029/2006JE002714.
Hansen, V.L., Willis, J.J., 1996. Structural analysis of a sampling of tesserae: implications for Venus geodynamics. Icarus 123, 296–312.
Hansen, V.L., Willis, J.J., Banerdt, W.B., 1997. Tectonic overview and synthesis In:
Bougher, S.W., Hunten, D.M., Phillips, R.J. (Eds.), Venus II Geology, Geophysics,
Atmosphere, and Solar Wind Environment. University Arizona Press, Tucson,
pp. 797–844.
Hansen, V.L., Willis, J.J., 1998. Ribbon terrain formation, southwestern fortuna
tessera, Venus: implications for lithosphere evolution. Icarus 132, 321–343.
Hansen, V.L., Phillips, R.G., Willis, J.J., Ghent, R.R., 2000. Structures in tessera terrain,
Venus: issues and answers. J. Geophys. Res. 105, 4135–4152.
Hansen, V.L., Olive, A., 2010. Artemis, Venus: the largest tectonomagmatic feature in
the solar system? Geology 38, 467–470.
Head, J.W., 1990. Assemblages of geologic/morphologic units in the northern
hemisphere of Venus. Earth Moon Planets 50–51, 391–408.
Head, J.W., Crumpler, L.S., Aubele, J.C., Guest, J.E., Saunders, R.S., 1992. Venus volcanism: classification of volcanic features and structures, associations, and
global distribution from Magellan data. J. Geophys. Res. 97, 13153–13197.
Head, J.W., Parmentier, E.M., Hess, P.C., 1994. Venus: vertical accretion of crust and
depleted mantle and implications for geological history and processes. Planet.
Space Sci. 42, 803–811.
Head, J.W., Ivanov, M.A., 1996. Tessera terrain formation and evolution on Venus:
evidence for phase III epeirogenic uplift in Ovda Regio. In: Proceedings of the
27th Lunar and Planetary Science Conference, pp. 517–518.
Head, J.W., Chapman, C.R., Strom, R.G., et al., 2011. Flood volcanism in the northern
high latitudes of Mercury revealed by MESSENGER. Science 333, 1853–1856.
http://dx.doi.org/10.1126/science.1211997.
Herrick, R.R., 1999. Small mantle upwellings are pervasive on Venus and Earth.
Geophys. Res. Lett. 26, 803–806.
Herrick, R.R., Sharpton, V.L., Malin, M.C., Lyons, S.N., Feely, K., 1997. Morphology and
morphometry of impact craters In: Bougher, S.W., Hunten, D.M., Phillips, R.J.
(Eds.), Venus II Geology, Geophysics, Atmosphere, and Solar Wind Environment. University Arizona Press, Tucson, pp. 1015–1046.
Ivanov, M.A., Head, J.W., 1996. Tessera terrain on Venus: a survey of the global
distribution, characteristics, and relation to surrounding units from Magellan
data. J. Geophys. Res. 101, 14861–14908.
Ivanov, M.A., Head, J.W., 2000. Stratigraphy of ridges and ribbons in tessera terrain,
Venus. In: Proceedings of the 31st Lunar and Planetary Science Conference,
Abstract 1233.
Ivanov, M.A., Head, J.W., 2001a. Geology of Venus: mapping of a global geotraverse
at 30 N latitude. J. Geophys. Res. 106, 17515–17566.
Ivanov, M.A., Head, J.W., 2001b. Geologic Map of the Lavinia Planitia Quadrangle (V55), Venus. USGS Scientific Investigations Map 2684.
Ivanov, M.A., Head, J.W., 2004a. Geologic map of the Atalanta Planitia (V4), Venus.
USGS Scientific Investigations Map 2792.
Ivanov, M.A., Head, J.W., 2004b. Stratigraphy of small shield volcanoes on Venus:
criteria for determining stratigraphic relationships and assessment of relative
age and temporal abundance. J. Geophys. Res. 109, E10001. http://dx.doi.org/
10.1029/2004JE002252.
Ivanov, M.A., Head, J.W., 2005a. Geologic map of the Nemesis Tesserae Quadrangle
(V-13), Venus. USGS Scientific Investigations Map 2870.
Ivanov, M.A., Head, J.W., 2006. Testing directional (evolutionary) and non-directional models of the geologic history of Venus: results of mapping in a
31
geotraverse along the equator of Venus. In: Proceedings of the 37th Lunar and
Planetary Science Conference, Abstract 1366.
Ivanov, M.A., Head, J.W., 2007. Semuni Dorsa: an extinct zone of convergence,
underthrusting and subduction on Venus? In: Proceedings of the 38th Lunar
and Planetary Science Conference, Abstract 1031.
Ivanov, M.A., Head, J.W., 2008a. Geologic Map of the Meskhent Tessera Quadrangle
Ivanov, M.A., Head, J.W., 2008b. Formation and evolution of Lakshmi Planum,
Venus: assessment of models using observations from geological mapping.
Planet. Space Sci. 56, 1949–1966.
Ivanov, M.A., Head, J.W., 2005b. Geologic Map of the Lakshmi Planum Quadrangle
Ivanov, M.A., Korteniemi, J., Kostama, V.-P., Aitolla, M., Raitala, J., Glamoclija, M.,
Marinangeli, L., Neukum, G., 2005. Major episodes of the hydrologic history in
the region of Hesperia Planum, Mars. J. Geophys. Res. 110, E12S21. http://dx.doi.
org/10.1029/2005JE002420.
Ivanov, M.A., Head, J.W., 2010. Geologic map of the Lakshmi Planum Quadrangle
Ivanov, M.A., Head, J.W., 2011. Global geological map of Venus. Planet. Space Sci. 59,
1559–1600.
Ivanov, M.A., Head, J.W., 2012. Formation and evolution of the midlands of Venus:
geological features and structures, stratigraphic relationships and geologic
history of the Fredegonde area (V-57). Planet. Space Sci. 73, 191–213.
Ivanov, M.A., Head, J.W., 2013. The history of volcanism on Venus. Planet. Space Sci.
84, 66–92.
Ivanov, M.A., Head, J.W., 2015. Volcanically embayed craters on Venus: testing the
catastrophic and equilibrium resurfacing models. Planet. Space Sci. 106, 116–121.
Jellinek, A.M., Lenardic, A., Manga, M., 2002. The influence of interior mantle
temperature on the structure of plumes: heads for Venus, tails for the Earth.
Geophys. Res. Lett. 29, 27-1–27-4. http://dx.doi.org/10.1029/2001GL014624.
Ji, S., Saruwatari, K., 1998. A revised model for the relationships between joint
spacing and layer thickness. J. Struct. Geol. 20, 1495–1508.
Klimczak, C., Watters, T.R., Ernst, C.M., et al., 2012. Deformation associated with
ghost craters and basins in volcanic smooth plains on Mercury: strain analysis
and implications for plains evolution. J. Geophys. Res. 117, E00L03. http://dx.
doi.org/10.1029/2012JE004100.
Koenig, E., Aydin, A., 1998. Evidence for large-scale strike–slip faulting on Venus.
Geology 26, 551–554.
Kreslavsky, M.A., Head, J.W., 1999. Morphometry of small shield volcanoes on
Venus: implications for the thickness of regional plains. J. Geophys. Res. 104,
18925–18932.
Kreslavsky, M.A., Ivanov, M.A., Head, J.W., 2015. The resurfacing history of Venus:
constraints from buffered crater densities. Icarus 250, 438–450.
Kryuchkov, V.P., 1990. Ridge belts: are they compressional or extensional structures? Earth Moon Planets 50–51, 471–491.
Kruychkov, V.P., 1992. Ridge belts on plains In: Barsukov, V.L., Basilevsky, A.T.,
pp. 96–112.
Kumar, P.S., 2005. J. Geophys. Res. 110, E07001. http://dx.doi.org/10.1029/
2004JE002387.
Ladeira, F.L., Price, N.J., 1981. Relationships between fracture spacing and bed
thickness. J. Struct. Geol. 3, 179–183.
Mastrapa, R., 1997. Thermal evolution of Venus: a preliminary study based on
tectonic feature spacing. In: Proceedings of the 28th Lunar and Planetary Science Conference, Abstract 1756.
Masursky, H., Eliason, E., E., Ford, P.G., et al., 1980. Pioneer-Venus radar results:
geology from the images and altimetry. J. Geophys. Res. 85, 8232–8260.
McGill, G.E., Campbell, B.A., 2006. Radar properties as clues to relative ages of ridge
belts and plains on Venus. J. Geophys. Res. 111, E12006. http://dx.doi.org/
10.1029/2006JE002705.
McGill, G.E., Steenstrup, S.J., Berton, C., Ford, P.G., 1981. Continental rifting and the
origin of Beta Regio, Venus. Geophys. Res. Lett. 8, 737–740.
McGill, G.E., Stofan, E.R., Smrekar, S.E., 2012. Venus tectonics In: Watters, T.R.,
Schultz, R.A. (Eds.), Planetary Tectonics. Cambridge University Press, Cambridge, pp. 81–121.
McGill, G.E., 2004. Geologic Map of the Bereghinya Planitia Quadrangle (V-8),
McGovern, P.J., Rumpf, M.E., Zimbelman, J.R., 2013. The influence of lithospheric
flexure on magma ascent at large volcanoes on Venus. J. Geophys. Res. 118,
2423–2437. http://dx.doi.org/10.1002/2013JE004455.
Montesi, L.G.J., Zuber, M.T., 2003. Clues to the lithospheric structure of Mars from
wrinkle ridge sets and localization instability. J. Geophys. Res. 108, 5048. http:
//dx.doi.org/10.1029/2002JE001974.
Moresi, L., Solomatov, V., 1998. Mantle convection with a brittle lithosphere:
thoughts on the global tectonic style of the Earth and Venus. Geophys. J. Int.
133, 669–682.
Moroz, V.I., 1983. Summary of preliminary results of the Venera 13 and Venera 14
missions In: Hunten, D.M., Colin, L., Donahue, T.M., Moroz, V.I. (Eds.), Venus.
University Arizona Press, Tucson, pp. 45–68.
Mueller, K., Golombek, M., 2004. Compressional structures on Mars. Annu. Rev.
Earth Planet. Sci. 32 435–464.
Musser, G.S., Squyres, S.W., 1997. A coupled thermal-mechanical model for corona
formation on Venus. J. Geophys. Res. 102, 6581–6595.
Parker, T.J., Saunders, R.S., 1994. Formation of complex ridged terrain through
progressive compressional deformation in the northetn Ovda Regio of Venus.
32
In: Proceedings of the 25th Lunar and Planetary Science Conference, pp. 1055–
1056.
Parmentier, E.M., Hess, P.C., 1992. Chemical differentiation of a convecting planetary interior: consequences for a one plate planet such as Venus. Geophys. Res.
Lett. 19, 2015–2018.
Pettengill, G.H., Eliason, E., Ford, P.G., et al., 1980. The surface of Venus. J. Geophys.
Res. 85, 8261–8270.
Phillips, R.J., Hansen, V.L., 1998. Geological evolution of Venus: rises, plains, plumes,
and plateaus. Science 279, 1492–1495.
Phillips, R.J., Grimm, R.E., Malin, M.C., 1991. Hot spot evolution and the global
tectonics of Venus. Science 252, 651–658.
Plescia, J.B., 1993. Wrinkle ridges of Arcadia Planitia, Mars. J. Geophys. Res. 98,
15049–15059.
Pritchard, M.E., Hansen, V.L., Willis, J.J., 1997. Structural evolution of western Fortuna Tessera, Venus. Geophys. Res. Lett. 24, 2339–2342.
Pronin, A.A., 1992. The Lakshmi phenomenon In: Barsukov, V.L., Basilevsky, A.T.,
pp. 68–81.
Pronin, A.A., Stofan, E.R., 1990. Coronae on Venus: morphology and distribution.
Icarus 87, 452–474.
Rosenberg, E., McGill, G.E., 2001. Geologic Map of the Pandrosos Dorsa Quadrangle
(V-5), Venus USGS Scientific Investigations Map 2721.
Schaber, G.G., Kirk, R.L., Strom, R.G., 1998. Data Base of Impact Craters on Venus
Based on Analysis of Magellan Radar Images and Altimetry Data. USGS Open
File Report 98-104.
Seif, A., 1983. Thermal structure of the atmosphere of Venus In: Hunten, D.M., Colin,
L., Donahue, T.M., Moroz, V.I. (Eds.), Venus. University Arizona Press, Tucson,
pp. 215–279.
Shalygin, E.V., Basilevsky, A.T., Markiewicz, W.J., Titov, D.V., Kreslavsky, M.A.,
Roatsch, Th, 2012. Search for ongoing volcanic activity on Venus: case study of
Maat Mons, Sapas Mons and Ozza Mons volcanoes. Planet. Space Sci. 73,
294–301.
Simons, M., Hager, B.H., Solomon, S.C., 1994. Global variations in the geoid/topography admittance of Venus. Science 264, 798–803.
Sjogren, W.L., Bills, B.G., Birkeland, P.W., et al., 1983. Venus gravity anomalies and
their correlations with topography. J. Geophys. Res. 88, 1119–1128.
Smrekar, S.E., Phillips, R.J., 1991. Venusian highlands: geoid to topography ratios
and their implications. Earth Planet. Sci. Lett. 107, 582–597.
Smrekar, S.E., Parmentier, M.E., 1996. The interaction of mantle plumes with surface thermal and chemical boundary layers: applications to hotspots on Venus.
J. Geophys. Res. 101, 5397–5410.
Smrekar, S.E., Stofan, E.R., 1997. Corona formation and heat loss on Venus by
coupled upwelling and delamination. Science 277, 1289–1294.
Smrekar, S.E., Kiefer, W.S., Stofan, E.R., 1997. Large volcanic rises on Venus In:
Bougher, S.W., Hunten, D.M., Phillips, R.J. (Eds.), Venus II Geology, Geophysics,
Atmosphere, and Solar Wind Environment. University Arizona Press, Tucson,
pp. 845–878.
Smrekar, S.E., Stofan, E.R., 1999. Origin of corona-dominated topographic rises on
Venus. Icarus 139, 100–115.
Solomatov, S.V., Moresi, L.-N., 1996. Stagnant lid convection on Venus. J. Geophys.
Res. 101, 4737–4753.
Solomon, S.C., 1993. A tectonic resurfacing model for Venus. In: Proceedings of the
24th Lunar and Planetary Science Conference, pp. 1331–1332.
Solomon, S.C., Head, J.W., 1982. Mechanisms for lithospheric heat transport on
Venus: implications for tectonic style and volcanism. J. Geophys. Res. 87,
9236–9246.
Solomon, S.C., Smrekar, S.E., Bindschadler, D.L., et al., 1992. Venus tectonics: an
overview of Magellan observations. J. Geophys. Res. 97, 13199–13255.
Squyres, S.W., Janes, D.M., Baer, G., Bindschadler, D.L., Schubert, G., Sharpton, V.L.,
Stofan, E.R., 1992. The morphology and evolution of coronae on Venus. J.
Geophys. Res. 97, 13611–13634.
Stofan, E.R., Guest, J., 2003. Geologic Map of the Aino Planitia Quadrangle (V-46),
Venus USGS Scientific Investigations Map 2779.
Stofan, E.R., Head, J.W., 1990. Coronae of Mnemosyne Regio: morphology and origin. Icarus 83, 216–243.
Stofan, E.R., Head, J.W., Campbell, D.B., Zisk, S.H., Bogomolov, A.F., Rzhiga, O.N.,
Basilevsky, A.T., Armand, N., 1989. Geology of a rift zone on Venus: Beta Regio
and Devana Chasma. Geol. Soc. Am. Bull. 101, 143–156.
Stofan, E.R., Bindschadler, D.L., Head, J.W., Parmentier, E.M., 1991. Corona structures
on Venus: models of origin. J. Geophys. Res. 96, 20933–20946.
Stofan, E.R., Sharpton, V.L., Schubert, G., Baer, G., Bindschadler, D.L., Janes, D.M.,
Squyres, S.W., 1992. Global distribution and characteristics of coronae and
related features on Venus: implications for origin and relation to mantle processes. J. Geophys. Res. 97, 13347–13378.
Stofan, E.R., Smrekar, S.E., Tapper, S.W., Guest, J.E., Grindrod, P.M., 2001. Preliminary
analysis of an expanded corona database for Venus. Geophys. Res. Lett. 28,
4267–4270.
Studd, D., Ernst, R.E., Samson, C., Grosfils, E.B., Head, J.W., Ivanov, M.A., 2010.
Radiating graben-fissure systems in Ulfrun Regio: a contribution to the Venus
global dyke swarm map. In: Proceedings of the 41st Lunar and Planetary Science Conference, Abstract 1950.
Studd, D., Ernst, R.E., Samson, C., 2011. Radiating graben-fissure systems in the
Ulfrun Regio area, Venus. Icarus 215, 279–291.
Sukhanov, A.L., 1992. Tesserae In: Barsukov, V.L., Basilevsky, A.T., Volkov, V.P.,
Zharkov, V.N. (Eds.), Venus Geology, Geochemistry, and Geophysics (Research
Results from the USSR). University of Arizona Press, Tucson, London, pp. 82–95.
Taylor, F.W., 2014. The Scientific Exploration of Venus. Cambridge University Press,
New York, p. 295.
Tuckwell, G.W., Ghail, R.C., 2003. A 400-km-scale strike-slip zone near the
boundary of Thetis Regio, Venus. Earth Planet. Sci. Lett. 211, 45–55.
van der Pluijm, B., Marshak, S., 2004. Earth Structure. An Introduction to Structural
Geology and Tectonics, 2nd edition. WW Norton & Company, New York, p. 672.
Vorder Bruegge, R.W., Head, J.W., 1991. Process of formation and evolution of
mountain belts on Venus. Geology 19, 885–888.
Watters, T.R., Schultz, R.A., Robinson, M.S., 2000. Displacement-length relations of
thrust faults associated with lobate scarps on Mercury and Mars: comparison
with terrestrial faults. Geophys. Res. Lett. 27, 3659–3662.
Wilhelms, D.E., 1974. Geologic mapping of the second planet, Pt.1: rationale and
general methods of Lunar geologic mapping In: Greeley, R., Schultz, P. (Eds.), A
Primer in Lunar Geology. Ames Research Center NASA, pp. 199–215.
Wilhelms, D.E., 1990. Geologic mapping In: Greeley, R., Batson, R.M. (Eds.), Planetary Mapping. Cambridge University Press, New York, Port Chester, Melbourne,
Sydney, pp. 208–260.
Wu, H., Pollard, D.D., 1995. An experimental study of the relationships between
joint spacing and layer thickness. J. Struct. Geol. 17, 887–905.
Young, D.A., Hansen, V.L., 2005. Poludnitsa Dorsa, Venus: history and context of a
deformation belt. J. Geophys. Res. 110, E03001. http://dx.doi.org/10.1029/
2004JE002280.

4574 - Ivanov, M. A., and J. W. Head III

Transcription

Similar documents

Presentation 2

by Gerard P. Kuiper June 28, 1969 Hypotheses on the composition

Sanchin and Tensho Masterclass.cdr - IOGKF

Landforms - rs181.k12.sd.us

West Texas Assemblies of God

Horse-hair Bridles Link History

3285 - Basilevsky, A. T., and J. W. Head

July 1973 - Warren Astronomical Society

to and print Sun Fun Karate Tournament