# Volcanic and Tectonic Constraints on the Evolution of Venus

Richard Ghail, Suzanne Smrekar, Thomas Widemann, Paul Byrne, Anna Gülcher, Joseph O'rourke, Madison Borrelli, Martha Gilmore, Robert Herrick, Mikhail Ivanov, et al.

► To cite this version:

Richard Ghail, Suzanne Smrekar, Thomas Widemann, Paul Byrne, Anna Gülcher, et al. Volcanic and Tectonic Constraints on the Evolution of Venus. Space Science Reviews, 2024, 220 (4), pp.36-10.1007/s11214-024-01065-2. http://dx.doi.org/10.1007/s11214-024-01065-2

HAL Id: hal-04589829

https://u-paris.hal.science/hal-04589829

Submitted on 27 May 2024

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
# Volcanic and Tectonic Constraints on the Evolution of Venus

Richard Ghail, Suzanne Smrekar, Thomas Widemann, Paul Byrne, Anna Gülcher, Joseph O'rourke, Madison Borrelli, Martha Gilmore, Robert Herrick, Mikhail Ivanov, et al.

► To cite this version:

Richard Ghail, Suzanne Smrekar, Thomas Widemann, Paul Byrne, Anna Gülcher, et al. Volcanic and Tectonic Constraints on the Evolution of Venus. Space Science Reviews, 2024, 220 (4), pp.36-10.1007/s11214-024-01065-2. http://dx.doi.org/10.1007/s11214-024-01065-2

HAL Id: hal-04589829

https://u-paris.hal.science/hal-04589829v1

Submitted on 27 May 2024

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
# Volcanic and Tectonic Constraints on the Evolution of Venus

Richard C. Ghail¹ ID • Suzanne E. Smrekar² ID • Thomas Widemann³,⁴ ID •

Paul K. Byrne⁵ ID • Anna J.P. Gülcher² ID • Joseph G. O'Rourke⁶ ID •

Madison E. Borrelli⁶ ID • Martha S. Gilmore⁷ ID • Robert R. Herrick⁸ ID •

Mikhail A. Ivanov⁹ ID • Ana-Catalina Plesa¹⁰ ID • Tobias Rolf¹¹ ID • Leah Sabbath² ID •

Joe W. Schools¹² ID • J. Gregory Shellnutt¹³ ID

Received: 6 March 2023 / Accepted: 27 March 2024
© The Author(s) 2024

## Abstract

Surface geologic features form a detailed record of Venus' evolution. Venus displays a pro-fusion of volcanic and tectonic features, including both familiar and exotic forms. One challenge to assessing the role of these features in Venus' evolution is that there are too few impact craters to permit age dates for specific features or regions. Similarly, without surface water, erosion is limited and cannot be used to evaluate age. These same observations indicate Venus has, on average, a very young surface (150–1000 Ma), with the most recent surface deformation and volcanism largely preserved on the surface except where covered by limited impact ejecta. In contrast, most geologic activity on Mars, the Moon, and Mercury occurred in the 1st billion years. Earth's geologic processes are almost all a result of plate tectonics. Venus' lacks such a network of connected, large scale plates, leaving the nature of Venus' dominant geodynamic process up for debate. In this review article, we describe Venus' key volcanic and tectonic features, models for their origin, and possible links to evolution. We also present current knowledge of the composition and thickness of the crust, lithospheric thickness, and heat flow given their critical role in shaping surface geology and interior evolution. Given Venus' hot lithosphere, abundant activity and potential analogues of continents, roll-back subduction, and microplates, it may provide insights into early Earth, prior to the onset of true plate tectonics. We explore similarities and differences between Venus and the Proterozoic or Archean Earth. Finally, we describe the future measurements needed to advance our understanding of volcanism, tectonism, and the evolution of Venus.

**Keywords** Venus • Formation of volcanic and tectonic features • Mantle dynamics • Subduction and plate tectonics • Topography • Hotspot volcanism • Plume-induced subduction • Crust and surface composition • Lithospheric thickness • Coronae • Lava channels • Venus resurfacing history
# Volcanic and Tectonic Constraints on the Evolution of Venus

Richard C. Ghail¹ ID • Suzanne E. Smrekar² ID • Thomas Widemann³,⁴ ID •

Paul K. Byrne⁵ ID • Anna J.P. Gülcher² ID • Joseph G. O'Rourke⁶ ID •

Madison E. Borrelli⁶ ID • Martha S. Gilmore⁷ ID • Robert R. Herrick⁸ ID •

Mikhail A. Ivanov⁹ ID • Ana-Catalina Plesa¹⁰ ID • Tobias Rolf¹¹ ID • Leah Sabbath² ID •

Joe W. Schools¹² ID • J. Gregory Shellnutt¹³ ID

Received: 6 March 2023 / Accepted: 27 March 2024

© The Author(s) 2024

## Abstract

Surface geologic features form a detailed record of Venus' evolution. Venus displays a pro-fusion of volcanic and tectonic features, including both familiar and exotic forms. One challenge to assessing the role of these features in Venus' evolution is that there are too few impact craters to permit age dates for specific features or regions. Similarly, without surface water, erosion is limited and cannot be used to evaluate age. These same observations indicate Venus has, on average, a very young surface (150–1000 Ma), with the most recent surface deformation and volcanism largely preserved on the surface except where covered by limited impact ejecta. In contrast, most geologic activity on Mars, the Moon, and Mercury occurred in the 1st billion years. Earth's geologic processes are almost all a result of plate tectonics. Venus' lacks such a network of connected, large scale plates, leaving the nature of Venus' dominant geodynamic process up for debate. In this review article, we describe Venus' key volcanic and tectonic features, models for their origin, and possible links to evolution. We also present current knowledge of the composition and thickness of the crust, lithospheric thickness, and heat flow given their critical role in shaping surface geology and interior evolution. Given Venus' hot lithosphere, abundant activity and potential analogues of continents, roll-back subduction, and microplates, it may provide insights into early Earth, prior to the onset of true plate tectonics. We explore similarities and differences between Venus and the Proterozoic or Archean Earth. Finally, we describe the future measurements needed to advance our understanding of volcanism, tectonism, and the evolution of Venus.

**Keywords** Venus • Formation of volcanic and tectonic features • Mantle dynamics • Subduction and plate tectonics • Topography • Hotspot volcanism • Plume-induced subduction • Crust and surface composition • Lithospheric thickness • Coronae • Lava channels • Venus resurfacing history

Venus: Evolution Through Time

Edited by Colin F. Wilson, Doris Breuer, Cédric Gillmann, Suzanne E. Smrekar, Tilman Spohn and Thomas Widemann

Extended author information available on the last page of the article
# 1 Introduction

This article describes current thinking on Venus' tectonic and volcanic processes, which provide essential insights on Venus' evolution through time. Given the often-noted gross similarities between Earth and Venus, an essential comparative planetology question is to know which processes differ or are the same between the two planets, and why? As the role of plate tectonics in Earth's long-term habitability becomes more evident, many exoplanet models have focussed on predicting the likelihood of this process. As the only other Earth-sized planet in the Solar System, Venus is the logical planet to explore this question (see Rolf et al. 2022 and many references therein). Venera measurements of radiogenic elements in Venus' crust show Earth-like concentrations, implying a similar internal heat budget. However, plate tectonics has not always been the dominant process on Earth either. Exactly when plate tectonics started and continents formed is debated, but certainly processes in the Archean (much of the first half of Earth's history) appear to have been very different to the modern Earth. The high surface temperature on Venus creates lithospheric temperatures similar to the elevated heat flow in the Archean (e.g., Van Kranendonk, 2010; Harris and Bédard 2014, 2015; Gerya et al. 2015), although with a transition to eclogite at a shallower depth than in Earth's present-day crust (e.g., Namiki and Solomon 1993). Much of Earth's early record has been erased by both plate tectonics and erosion. Tectonic and volcanic processes on Venus today may therefore inform our understanding of Archean Earth and vice versa. Some features on Venus appear consistent with localized subduction of the lithosphere (e.g., Schubert and Sandwell 1995; Harris and Bédard 2014, 2015). As subduction is believed to be the first step in creating a plate tectonic system, understanding the conditions necessary to permit this process to occur, is the absence of a plate tectonic framework. Other key issues include the formation of Venus' tesserae and whether or not Earth's Archean cratons formed in an analogous manner.

A key constraint on interpreting geologic processes on another planet is the size and spatial distribution of impact craters. Venus has only ~1000 impact craters and their distribution cannot be distinguished from a random one. The average surface age is very young, relative to other terrestrial planets, but poorly constrained (estimates span at least ~0.15–1 Ga). The processes that removed older terrains thereby removing craters are controversial (Herrick et al. 2023, this collection). The absence of large regions of conspicuously different ages on Venus has been used to infer a very different tectonic style from Earth (e.g., Herrick et al. 2023, this collection). Furthermore, only ~10% of the craters are widely regarded as modified by subsequent geologic activity. However, many more craters might have suffered modification by sedimentary (Carter et al. 2023, this collection) and/or volcanic processes (e.g., Herrick and Rumpf 2011; Herrick et al. 2023, this collection). The random spatial distribution and ambiguous modification state of the impact craters has led to competing interpretations of Venus' impact record. In the “catastrophic resurfacing” model, the entire surface was wiped clean of impact craters, either by burial under a thick layer of lavas (e.g., Schaber et al. 1992; Head et al. 1994; Strom et al. 1994) and/or by recycling of the surface via overturn of the lithosphere (e.g., Parmentier and Hess 1992). Subsequently, geologic activity declined, allowing the surface to accumulate impact craters at random. The globally ancient surfaces of Mars, Mercury, and the Moon are in a stagnant lid convective regime, with interior heat lost primarily through conduction across a thick, immobile lithosphere. However, for Venus, a fully stagnant lid does not provide sufficient heat loss, leading to the idea that Venus has cycled between a stagnant lid regime today and a past plate tectonic regime, allowing for greater heat loss over time (Turcotte et al. 1999), consistent with the inference of a past catastrophic resurfacing event.
# 1 Introduction

This article describes current thinking on Venus' tectonic and volcanic processes, which provide essential insights on Venus' evolution through time. Given the often-noted gross similarities between Earth and Venus, an essential comparative planetology question is to know which processes differ or are the same between the two planets, and why? As the role of plate tectonics in Earth's long-term habitability becomes more evident, many exoplanet models have focussed on predicting the likelihood of this process. As the only other Earth-sized planet in the Solar System, Venus is the logical planet to explore this question (see Rolf et al. 2022 and many references therein). Venera measurements of radiogenic elements in Venus' crust show Earth-like concentrations, implying a similar internal heat budget. However, plate tectonics has not always been the dominant process on Earth either. Exactly when plate tectonics started and continents formed is debated, but certainly processes in the Archean (much of the first half of Earth's history) appear to have been very different to the modern Earth. The high surface temperature on Venus creates lithospheric temperatures similar to the elevated heat flow in the Archean (e.g., Van Kranendonk, 2010; Harris and Bédard 2014, 2015; Gerya et al. 2015), although with a transition to eclogite at a shallower depth than in Earth's present-day crust (e.g., Namiki and Solomon 1993). Much of Earth's early record has been erased by both plate tectonics and erosion. Tectonic and volcanic processes on Venus today may therefore inform our understanding of Archean Earth and vice versa. Some features on Venus appear consistent with localized subduction of the lithosphere (e.g., Schubert and Sandwell 1995; Harris and Bédard 2014, 2015). As subduction is believed to be the first step in creating a plate tectonic system, understanding the conditions necessary to permit this process to occur, is the absence of a plate tectonic framework. Other key issues include the formation of Venus' tesserae and whether or not Earth's Archean cratons formed in an analogous manner.

A key constraint on interpreting geologic processes on another planet is the size and spatial distribution of impact craters. Venus has only ~1000 impact craters and their distribution cannot be distinguished from a random one. The average surface age is very young, relative to other terrestrial planets, but poorly constrained (estimates span at least ~0.15–1 Ga). The processes that removed older terrains thereby removing craters are controversial (Herrick et al. 2023, this collection). The absence of large regions of conspicuously different ages on Venus has been used to infer a very different tectonic style from Earth (e.g., Herrick et al. 2023, this collection). Furthermore, only ~10% of the craters are widely regarded as modified by subsequent geologic activity. However, many more craters might have suffered modification by sedimentary (Carter et al. 2023, this collection) and/or volcanic processes (e.g., Herrick and Rumpf 2011; Herrick et al. 2023, this collection). The random spatial distribution and ambiguous modification state of the impact craters has led to competing interpretations of Venus' impact record. In the “catastrophic resurfacing” model, the entire surface was wiped clean of impact craters, either by burial under a thick layer of lavas (e.g., Schaber et al. 1992; Head et al. 1994; Strom et al. 1994) and/or by recycling of the surface via overturn of the lithosphere (e.g., Parmentier and Hess 1992). Subsequently, geologic activity declined, allowing the surface to accumulate impact craters at random. The globally ancient surfaces of Mars, Mercury, and the Moon are in a stagnant lid convective regime, with interior heat lost primarily through conduction across a thick, immobile lithosphere. However, for Venus, a fully stagnant lid does not provide sufficient heat loss, leading to the idea that Venus has cycled between a stagnant lid regime today and a past plate tectonic regime, allowing for greater heat loss over time (Turcotte et al. 1999), consistent with the inference of a past catastrophic resurfacing event.
The opposite end-member is the ‘equilibrium resurfacing’ hypothesis (e.g., Phillips et al. 1992). Under this scenario, impact craters are removed at a roughly constant rate by geo-logic activity that has no spatial organization, in contrast to Earth. Perhaps more likely is an intermediate hypothesis of ‘regional equilibrium resurfacing’ that allows for spatial and temporal variability of regions less than ~1000 km in diameter, consistent with the impact crater distribution (e.g., Phillips et al. 1992). Herrick et al. (2023, this collection) provides a thorough discussion of these models and their implications. We introduce these interpre-tations here because of the fundamental role that both the mean surface age and local vari-ability play in framing the geologic evolution of Venus and the interpretation of interior and atmospheric processes (e.g., Rolf et al. 2022, this collection). Recently, the concept of a par-tially mobile ‘squishy’ lid regime has been proposed, in which intrusive volcanism allows for plate-tectonic-like levels of heat loss (Lourenço et al. 2020). Below, we discuss a range of geologic features observed on Venus, along with implications for lithospheric structure and convective regime.

In parallel with the notion of ‘episodic vs. equilibrium’ resurfacing is the concept of ‘directional vs non-directional’ geologic history (Guest and Stofan 1999). Does the variety of geologic features observed on Venus occur in a specific sequence everywhere, consis-tent with an episodic (or ‘catastrophic’) resurfacing event? Ivanov and Head (2011) mapped the whole Venus surface at 1:10 million scale and identified a global stratigraphic sequence with 13 distinct members, with tesserae identified everywhere as the oldest unit, then various deformed plains, ‘groove’ (extensional) and mountain (compressional) belts, through unde-formed plains, to ‘lobate’ (lava flow) plains and rift zones. The authors argued that since this sequence is never found next to, and offset from, an adjacent sequence, the sequence must represent a global change through time, i.e., a directional geologic history. Arguments against this sequence (Guest and Stofan 1999) take two principal forms: first, that the strati-graphic methods and interpretation of some authors (e.g., Ivanov and Head 2011) mix mate-rial (rock) units with tectonic (structural) events; and second, that radar is sensitive to surface texture and relatively insensitive to rock type. For example, the regional plains, which cover a third of the planet, may have very different ages but appear texturally the same, and hence cannot be used as a stratigraphic marker. A further problem for interpretation is that since radar is very sensitive to surface roughness near the radar wavelength, apparent boundaries and onlap relationships, such as the ‘flooding’ of tesserae by regional plains, may be a result of gradational downslope grain size changes, indicating the opposite age relationship (Carter et al. 2023, this collection, and references therein). Further, others argue that this sequence is not observed everywhere and that in places, tesserae may be younger than the plains (e.g., Ghail 2002; Ivanov and Head 1996; Gilmore and Head 2000). The scarcity of impact craters and the lack of any other independent dating technique hamper a resolution to this question.

Our understanding of current processes and conditions effectively provide boundary con-
ditions for extrapolation backward into Venus' geologic and convective history (see Rolf
et al. 2022, this collection). The size, type, distribution, and (where available) chemistry of
volcanic features offers a window into interior processes such as upwelling/downwelling,
decompression melting and mantle temperature. Venus hosts a wide array of volcanic fea-
tures, from those seen on Earth including volcanic edifices and lava flows, to those unique
(at least in abundance) to Venus, such as novae and coronae (e.g. Glaze et al. 2002; Krassil-
nikov and Head 2003). The overlap between volcanoes, coronae, and novae is an interesting
puzzle, as discussed below. Tectonic features provide constraints on the rheology and thick-
ness of the lithosphere and the origin and magnitude of driving stress fields. Gravity data
provide further information on the thickness of the crust, the lithosphere, and subsurface
density variations due to processes such as mantle upwellings. The low resolution of cur-
rent datasets often results in multiple interpretations of a given feature or process. The goal
The opposite end-member is the ‘equilibrium resurfacing’ hypothesis (e.g., Phillips et al. 1992). Under this scenario, impact craters are removed at a roughly constant rate by geo-logic activity that has no spatial organization, in contrast to Earth. Perhaps more likely is an intermediate hypothesis of ‘regional equilibrium resurfacing’ that allows for spatial and temporal variability of regions less than ~1000 km in diameter, consistent with the impact crater distribution (e.g., Phillips et al. 1992). Herrick et al. (2023, this collection) provides a thorough discussion of these models and their implications. We introduce these interpre-tations here because of the fundamental role that both the mean surface age and local vari-ability play in framing the geologic evolution of Venus and the interpretation of interior and atmospheric processes (e.g., Rolf et al. 2022, this collection). Recently, the concept of a par-tially mobile ‘squishy’ lid regime has been proposed, in which intrusive volcanism allows for plate-tectonic-like levels of heat loss (Lourenço et al. 2020). Below, we discuss a range of geologic features observed on Venus, along with implications for lithospheric structure and convective regime.

In parallel with the notion of ‘episodic vs. equilibrium’ resurfacing is the concept of ‘directional vs non-directional’ geologic history (Guest and Stofan 1999). Does the variety of geologic features observed on Venus occur in a specific sequence everywhere, consis-tent with an episodic (or ‘catastrophic’) resurfacing event? Ivanov and Head (2011) mapped the whole Venus surface at 1:10 million scale and identified a global stratigraphic sequence with 13 distinct members, with tesserae identified everywhere as the oldest unit, then various deformed plains, ‘groove’ (extensional) and mountain (compressional) belts, through unde-formed plains, to ‘lobate’ (lava flow) plains and rift zones. The authors argued that since this sequence is never found next to, and offset from, an adjacent sequence, the sequence must represent a global change through time, i.e., a directional geologic history. Arguments against this sequence (Guest and Stofan 1999) take two principal forms: first, that the strati-graphic methods and interpretation of some authors (e.g., Ivanov and Head 2011) mix mate-rial (rock) units with tectonic (structural) events; and second, that radar is sensitive to surface texture and relatively insensitive to rock type. For example, the regional plains, which cover a third of the planet, may have very different ages but appear texturally the same, and hence cannot be used as a stratigraphic marker. A further problem for interpretation is that since radar is very sensitive to surface roughness near the radar wavelength, apparent boundaries and onlap relationships, such as the ‘flooding’ of tesserae by regional plains, may be a result of gradational downslope grain size changes, indicating the opposite age relationship (Carter et al. 2023, this collection, and references therein). Further, others argue that this sequence is not observed everywhere and that in places, tesserae may be younger than the plains (e.g., Ghail 2002; Ivanov and Head 1996; Gilmore and Head 2000). The scarcity of impact craters and the lack of any other independent dating technique hamper a resolution to this question.

Our understanding of current processes and conditions effectively provide boundary con-
ditions for extrapolation backward into Venus' geologic and convective history (see Rolf
et al. 2022, this collection). The size, type, distribution, and (where available) chemistry of
volcanic features offers a window into interior processes such as upwelling/downwelling,
decompression melting and mantle temperature. Venus hosts a wide array of volcanic fea-
tures, from those seen on Earth including volcanic edifices and lava flows, to those unique
(at least in abundance) to Venus, such as novae and coronae (e.g. Glaze et al. 2002; Krassil-
nikov and Head 2003). The overlap between volcanoes, coronae, and novae is an interesting
puzzle, as discussed below. Tectonic features provide constraints on the rheology and thick-
ness of the lithosphere and the origin and magnitude of driving stress fields. Gravity data
provide further information on the thickness of the crust, the lithosphere, and subsurface
density variations due to processes such as mantle upwellings. The low resolution of cur-
rent datasets often results in multiple interpretations of a given feature or process. The goal
here is not to provide a single view, but to represent a range of possible interpretations, the constraints available from current data, and to outline specific needs for future datasets.

A rocky body's crustal layer provides a complex archive for the planet's evolution. The crust may be primary, a direct consequence of a magma ocean freezing, secondary, a result of active magmatism and volcanism in the solidified interior, or tertiary, which invokes remelting of primary and/or secondary crust (e.g., like the Earth's continental crust). In the case of Venus, crustal composition is dominantly basaltic (Hess and Head 1990) and its young surface age dictates that most of its volume is likely secondary crust. Tessera terrains may represent tertiary crust (see Gilmore et al., 2023, this collection). The formation of eclogite from basalt may play a key role on Venus in terms of both mantle dynamics and the formation of specific volcanic and tectonic features. Delamination driven by eclogite formation has been proposed to drive the formation of both coronae (e.g., Piskorz et al. 2014; Gülcher et al. 2023) and tessera plateaus (e.g., Bindschadler and Parmentier 1990). Delamination would also result in the formation of secondary or tertiary crust (Elkins-Tanton et al. 2007). Whether or not delamination leads to mobilisation of an otherwise stagnant lid depends on the details of the mantle convection models (e.g., Armann and Tackley 2012; Rolf et al. 2018; Adams et al. 2022, see Rolf et al. 2022, this collection). Mantle convection models are useful tools to simulate the growth history of Venus' crust through its long-term evolution and how this relates to the planet's thermal state and geodynamic regime. A stagnant-lid scenario may lead to a relatively hot mantle, promoting magmatism and crustal growth. In contrast, episodically or partially mobile (plutonic squishy lid) lithosphere may reset crustal thickness and therefore limit its effective thickness (see Rolf et al. 2022, this collection, for details).

## 2 Overview of Volcanic Features

The surface of Venus is replete with volcanoes—that is, discrete edifices, although there is evidence for non-constructional volcanoes such as fissure vents, too. Early post-Magellan surveys (e.g., Head et al. 1992) reported on the positional details of almost 1,700 volcanic edifices across the planet, finding that roughly half the large volcanoes on Venus are located in the Beta, Atla, and Themis region, which covers less than a quarter of the surface (Crumpler et al. 1993). The largest volcanoes are well in excess of 100 km across: some, such as Maat (Fig. 1), Sif, and Tepev Montes, are several hundred kilometres in diameter (e.g., Janle et al. 1988; Senske et al. 1992), and lava flows as long as ~500–1000 km are observed (Head et al. 1992). Indeed, vast lava flows are a common feature of the largest volcanoes on Venus, which are typically located atop topographic rises (e.g., Ivanov and Head 2013). These edifices are classic shield volcanoes, being far broader than they are tall; generally, Venus' largest volcanoes are up to 9 km tall, comparable to the largest volcanoes on Earth (Fig. 1).

Venus also exhibits ~200 large flow fields, with areal extents of ~39,000 to 744,000 km² (e.g., Magee and Head 2001). The majority of these flow fields emanate from either central volcanos or coronae, and are thus covered under the discussion of these features below. In addition, another ~40 large flow fields occur in association with chasmata or fracture zones (e.g., Magee and Head 2001; Lancaster et al. 1995). These tectonic features are discussed below as well.

Volcanoes less than 100 km in diameter are distributed widely across the planet and are often characterized by radial lava flows and conical or shield-like shapes; some have flat-topped summits (Head et al. 1992). By far the greatest number of individual edifices are
here is not to provide a single view, but to represent a range of possible interpretations, the constraints available from current data, and to outline specific needs for future datasets.

A rocky body's crustal layer provides a complex archive for the planet's evolution. The crust may be primary, a direct consequence of a magma ocean freezing, secondary, a result of active magmatism and volcanism in the solidified interior, or tertiary, which invokes remelting of primary and/or secondary crust (e.g., like the Earth's continental crust). In the case of Venus, crustal composition is dominantly basaltic (Hess and Head 1990) and its young surface age dictates that most of its volume is likely secondary crust. Tessera terrains may represent tertiary crust (see Gilmore et al., 2023, this collection). The formation of eclogite from basalt may play a key role on Venus in terms of both mantle dynamics and the formation of specific volcanic and tectonic features. Delamination driven by eclogite formation has been proposed to drive the formation of both coronae (e.g., Piskorz et al. 2014; Gülcher et al. 2023) and tessera plateaus (e.g., Bindschadler and Parmentier 1990). Delamination would also result in the formation of secondary or tertiary crust (Elkins-Tanton et al. 2007). Whether or not delamination leads to mobilisation of an otherwise stagnant lid depends on the details of the mantle convection models (e.g., Armann and Tackley 2012; Rolf et al. 2018; Adams et al. 2022, see Rolf et al. 2022, this collection). Mantle convection models are useful tools to simulate the growth history of Venus' crust through its long-term evolution and how this relates to the planet's thermal state and geodynamic regime. A stagnant-lid scenario may lead to a relatively hot mantle, promoting magmatism and crustal growth. In contrast, episodically or partially mobile (plutonic squishy lid) lithosphere may reset crustal thickness and therefore limit its effective thickness (see Rolf et al. 2022, this collection, for details).

## 2 Overview of Volcanic Features

The surface of Venus is replete with volcanoes—that is, discrete edifices, although there is evidence for non-constructional volcanoes such as fissure vents, too. Early post-Magellan surveys (e.g., Head et al. 1992) reported on the positional details of almost 1,700 volcanic edifices across the planet, finding that roughly half the large volcanoes on Venus are located in the Beta, Atla, and Themis region, which covers less than a quarter of the surface (Crumpler et al. 1993). The largest volcanoes are well in excess of 100 km across: some, such as Maat (Fig. 1), Sif, and Tepev Montes, are several hundred kilometres in diameter (e.g., Janle et al. 1988; Senske et al. 1992), and lava flows as long as ~500–1000 km are observed (Head et al. 1992). Indeed, vast lava flows are a common feature of the largest volcanoes on Venus, which are typically located atop topographic rises (e.g., Ivanov and Head 2013). These edifices are classic shield volcanoes, being far broader than they are tall; generally, Venus' largest volcanoes are up to 9 km tall, comparable to the largest volcanoes on Earth (Fig. 1).

Venus also exhibits ~200 large flow fields, with areal extents of ~39,000 to 744,000 km² (e.g., Magee and Head 2001). The majority of these flow fields emanate from either central volcanos or coronae, and are thus covered under the discussion of these features below. In addition, another ~40 large flow fields occur in association with chasmata or fracture zones (e.g., Magee and Head 2001; Lancaster et al. 1995). These tectonic features are discussed below as well.

Volcanoes less than 100 km in diameter are distributed widely across the planet and are often characterized by radial lava flows and conical or shield-like shapes; some have flat-topped summits (Head et al. 1992). By far the greatest number of individual edifices are
<figure><img src="image_2.png" /><figcaption>**Fig. 1** Maat Mons and Mauna Kea (Hawaii) both reach ~9 km of prominence, but Maat Mons is several hundred km across. For comparison, Olympus Mons (Mars) reaches ~24 km of prominence, but still is narrower than Maat Mons.</figcaption></figure>

those less than 20 km in diameter (e.g., Crumpler et al. 1997; Grindrod et al. 2010), with re-
cent mapping suggesting that there are at least tens of thousands of volcanoes <5 km across
(Hahn and Byrne, 2022). Crumpler et al. (1997), classified volcanic edifices as “large”, “in-
termediate” and “small” on the basis of Venera 15 and 16 results; large volcanoes are those
equal to or larger than 100 km, intermediate volcanoes are those ≥ 20 km and < 100 km,
and small volcanoes are those < 20 km in diameter, regardless of their geologic significance.
Establishing the exact shape of these small edifices is difficult with the limited topographic
resolution of Magellan and so determining whether they are conical or truly shield-like in ap-
pearance requires new data from future missions. Given their prevalence, volcanic processes
likely have a key role in forming the young surface through the burial of impact craters and
other old terrains (see Herrick et al. 2023, this collection). As on Earth, some volcanism can
be linked directly to decompression melting above mantle plumes, while other features may
be associated with lithospheric extension, delamination or dripping, or other processes (e.g.
Head et al. 1992; Martin et al. 2007). The features that can most clearly be linked to mantle
plumes are the ten or so volcanic rises, which share many key characteristics with Hawaii
and other terrestrial hotspots (Stofan and Smrekar 2005).

The ~500 enigmatic coronae (mean diameter ~250 km) are widely distributed across the surface (e.g., Stofan et al. 1992), although ~62% occur in association with rifts or fracture belts (e.g., Glaze et al. 2002). These features likely have multiple formation mechanisms, but many probably formed above small-scale mantle plumes (e.g., Gülcher et al. 2020). Some proposed mechanisms have terrestrial analogues, such as large igneous provinces (LIPs) and plume-induced subduction. However, these processes are not active on Earth today. Two key questions for the formation of coronae are: (i) why do coronae form on Venus but not on present-day Earth, and (ii) why does Venus appear to have two largely distinct scales of mantle upwellings? For the latter question, one hypothesis is that different volcanic features, such as volcanic rises and coronae, formed at different times in Venus' evolution; another is that layered (or otherwise variable scales of) mantle convection generates different features (see Rolf et al. 2022, this collection).

**2.1 Spatial Distribution of Volcanism**

Given the seeming lack of obvious fluvial or aeolian sedimentary processes (although see Carter et al. 2023, this collection), volcanism is the major contributor to the formation of surface material on Venus (Head 1990) at least during its observable part of geologic history.
<figure><img src="image_2.png" /><figcaption>**Fig. 1** Maat Mons and Mauna Kea (Hawaii) both reach ~9 km of prominence, but Maat Mons is several hundred km across. For comparison, Olympus Mons (Mars) reaches ~24 km of prominence, but still is narrower than Maat Mons.</figcaption></figure>

those less than 20 km in diameter (e.g., Crumpler et al. 1997; Grindrod et al. 2010), with re-
cent mapping suggesting that there are at least tens of thousands of volcanoes <5 km across
(Hahn and Byrne, 2022). Crumpler et al. (1997), classified volcanic edifices as “large”, “in-
termediate” and “small” on the basis of Venera 15 and 16 results; large volcanoes are those
equal to or larger than 100 km, intermediate volcanoes are those ≥ 20 km and < 100 km,
and small volcanoes are those < 20 km in diameter, regardless of their geologic significance.
Establishing the exact shape of these small edifices is difficult with the limited topographic
resolution of Magellan and so determining whether they are conical or truly shield-like in ap-
pearance requires new data from future missions. Given their prevalence, volcanic processes
likely have a key role in forming the young surface through the burial of impact craters and
other old terrains (see Herrick et al. 2023, this collection). As on Earth, some volcanism can
be linked directly to decompression melting above mantle plumes, while other features may
be associated with lithospheric extension, delamination or dripping, or other processes (e.g.
Head et al. 1992; Martin et al. 2007). The features that can most clearly be linked to mantle
plumes are the ten or so volcanic rises, which share many key characteristics with Hawaii
and other terrestrial hotspots (Stofan and Smrekar 2005).

The ~500 enigmatic coronae (mean diameter ~250 km) are widely distributed across the surface (e.g., Stofan et al. 1992), although ~62% occur in association with rifts or fracture belts (e.g., Glaze et al. 2002). These features likely have multiple formation mechanisms, but many probably formed above small-scale mantle plumes (e.g., Gülcher et al. 2020). Some proposed mechanisms have terrestrial analogues, such as large igneous provinces (LIPs) and plume-induced subduction. However, these processes are not active on Earth today. Two key questions for the formation of coronae are: (i) why do coronae form on Venus but not on present-day Earth, and (ii) why does Venus appear to have two largely distinct scales of mantle upwellings? For the latter question, one hypothesis is that different volcanic features, such as volcanic rises and coronae, formed at different times in Venus' evolution; another is that layered (or otherwise variable scales of) mantle convection generates different features (see Rolf et al. 2022, this collection).

## 2.1 Spatial Distribution of Volcanism

Given the seeming lack of obvious fluvial or aeolian sedimentary processes (although see Carter et al. 2023, this collection), volcanism is the major contributor to the formation of surface material on Venus (Head 1990) at least during its observable part of geologic history.
<figure><img src="image_3.png" /><figcaption>**Fig. 2** - Examples of volcanic edifices on Venus: (a) a cluster of small edifices (< 20 km in diameter) that populate an exposure of shield plains; (b) a cluster of steep-sided domes that represent the intermediate volcanoes (from Crumpler and Aubele 2000); (c) a large volcano (> 100 km) of Sapas Mons (from Crumpler and Aubele 2000)</figcaption></figure>

Indeed, Venus is a “volcanic” planet on which volcanic materials compose ~92% of the surface (e.g., Ivanov and Head 2011). The remaining surface consists of tessera (as discussed below in Sect. 3.1).

The most areally extensive volcanic landforms on Venus are volcanic plains, some of which are as large as a few million square kilometres (e.g., Ivanov and Head 2011, 2013).

Small edifices (Fig. 2a) populate the surface of shield plains and are the most abundant volcanic form (Addington 2001; Crumpler et al. 1993; Hahn and Byrne, 2022). The exact number of these volcanoes is unknown but it may exceed the tens of thousands. These edifices are relatively small with the decile diameter range (central 80% of the population) from ~2.5 to ~5.5 km (Hahn and Byrne, 2022). In an early global catalogue of volcanic landforms of Venus (Head et al. 1992), the sub-population of small volcanoes on the shield plains were assigned to a class of “shield fields,” which may represent isolated outcrops of shield plains embayed by younger materials (e.g., Ivanov and Head 2004).

A recently compiled, new global volcano catalogue for Venus indicates the presence of more than 85,000 edifices, the vast majority of which (~52,000) are barely resolvable with available Magellan radar image data (Hahn and Byrne, 2022) (Fig. 3a). Of those volcanoes that are more readily identifiable, ~32,500 are <5 km in diameter, 729 are 5–100 km in diameter, and 118 have diameters greater than 100 km. Previous studies had further grouped volcanoes into “fields” (e.g., Grosfils et al. 2000; Pavri et al. 1992); Hahn and Byrne (2022) found that (for specific criteria, including 25 edifices as a minimum number) the global volcano population could be grouped into 566 fields comprising about 65% of all mapped volcanoes; the remainder did not fall within clusters of ≥25 edifices.

The spatial distribution of volcanoes on Venus (Fig. 3a) appears to be drastically different from that on Earth, where the active volcanoes are mostly occurring within zones along convergent and divergent boundaries and along hotspot tracks (Fig. 3b) (Byrne and Krishnamoorthy 2022), although it should be noted that the distribution of (presumably) active volcanoes on Venus is currently unknown. Even so, the difference in spatial distribution suggests that Venus possesses a mechanism of the internal heat leakage which is dissimilar to terrestrial plate tectonics (e.g., Solomon and Head 1982; Solomon et al. 1992; Stofan and Saunders 1990). Grossly speaking, volcanoes on Venus show slightly greater spatial concentration in the Beta–Atla–Themis region and in a region north of Aphrodite Terra relative to the rest of the planet surface (Hahn & Byrne, 2022) (Fig. 3a).
<figure><img src="image_3.png" /><figcaption>**Fig. 2** - Examples of volcanic edifices on Venus: (a) a cluster of small edifices (< 20 km in diameter) that populate an exposure of shield plains; (b) a cluster of steep-sided domes that represent the intermediate volcanoes (from Crumpler and Aubele 2000); (c) a large volcano (> 100 km) of Sapas Mons (from Crumpler and Aubele 2000)</figcaption></figure>

Indeed, Venus is a “volcanic” planet on which volcanic materials compose ~92% of the surface (e.g., Ivanov and Head 2011). The remaining surface consists of tessera (as discussed below in Sect. 3.1).

The most areally extensive volcanic landforms on Venus are volcanic plains, some of which are as large as a few million square kilometres (e.g., Ivanov and Head 2011, 2013).

Small edifices (Fig. 2a) populate the surface of shield plains and are the most abundant volcanic form (Addington 2001; Crumpler et al. 1993; Hahn and Byrne, 2022). The exact number of these volcanoes is unknown but it may exceed the tens of thousands. These edifices are relatively small with the decile diameter range (central 80% of the population) from ~2.5 to ~5.5 km (Hahn and Byrne, 2022). In an early global catalogue of volcanic landforms of Venus (Head et al. 1992), the sub-population of small volcanoes on the shield plains were assigned to a class of “shield fields,” which may represent isolated outcrops of shield plains embayed by younger materials (e.g., Ivanov and Head 2004).

A recently compiled, new global volcano catalogue for Venus indicates the presence of more than 85,000 edifices, the vast majority of which (~52,000) are barely resolvable with available Magellan radar image data (Hahn and Byrne, 2022) (Fig. 3a). Of those volcanoes that are more readily identifiable, ~32,500 are <5 km in diameter, 729 are 5–100 km in diameter, and 118 have diameters greater than 100 km. Previous studies had further grouped volcanoes into “fields” (e.g., Grosfils et al. 2000; Pavri et al. 1992); Hahn and Byrne (2022) found that (for specific criteria, including 25 edifices as a minimum number) the global volcano population could be grouped into 566 fields comprising about 65% of all mapped volcanoes; the remainder did not fall within clusters of ≥25 edifices.

The spatial distribution of volcanoes on Venus (Fig. 3a) appears to be drastically different from that on Earth, where the active volcanoes are mostly occurring within zones along convergent and divergent boundaries and along hotspot tracks (Fig. 3b) (Byrne and Krishnamoorthy 2022), although it should be noted that the distribution of (presumably) active volcanoes on Venus is currently unknown. Even so, the difference in spatial distribution suggests that Venus possesses a mechanism of the internal heat leakage which is dissimilar to terrestrial plate tectonics (e.g., Solomon and Head 1982; Solomon et al. 1992; Stofan and Saunders 1990). Grossly speaking, volcanoes on Venus show slightly greater spatial concentration in the Beta–Atla–Themis region and in a region north of Aphrodite Terra relative to the rest of the planet surface (Hahn & Byrne, 2022) (Fig. 3a).
<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /><figcaption>**Fig. 3** The spatial distribution of volcanoes on: (a) Venus, and (b) Earth. The databases of Venusian volcanic edifices are from Hahn and Byrne (2022). Terrestrial volcanoes are those that are thought to have been active during the Holocene volcanoes, which are mostly located at plate boundaries (Global Volcanism Program 2023). This database from the Smithsonian also has a list of terrestrial volcanoes that may have been active during the Pleistocene</figcaption></figure>

## 2.1.1 Lava Channels

Lava channels are widespread on Venus. Almost 200 such features were mapped using Magellan data, divided into six different types (e.g., Komatsu et al. 1993): sinuous rilles; canali; simple channels with associated flow margins; complex channels with flow margins; com-
<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /><figcaption>**Fig. 3** The spatial distribution of volcanoes on: (a) Venus, and (b) Earth. The databases of Venusian volcanic edifices are from Hahn and Byrne (2022). Terrestrial volcanoes are those that are thought to have been active during the Holocene volcanoes, which are mostly located at plate boundaries (Global Volcanism Program 2023). This database from the Smithsonian also has a list of terrestrial volcanoes that may have been active during the Pleistocene</figcaption></figure>

### 2.1.1 Lava Channels

Lava channels are widespread on Venus. Almost 200 such features were mapped using Magellan data, divided into six different types (e.g., Komatsu et al. 1993): sinuous rilles; canali; simple channels with associated flow margins; complex channels with flow margins; com-
plex channels without flow margins; and compound channels. Some of these channels resemble features seen on Mars and Earth's Moon. For example, sinuous rilles have large depth-to-width ratios and morphologies similar to those of lunar and Martian rilles (e.g., Komatsu and Baker 1994; Oshigami et al. 2009). Usually, no lateral flow deposits are observed near rilles, but flow deposits are sometimes found at their termini. Channel sources are sometimes unidentifiable, but they are often found to originate in depressions. Mappers have identified ~59 rilles with average lengths of ~10–300 km and widths up to several kilometres (e.g., Baker et al. 1992; Komatsu et al. 1993). More than half of the rilles are associated with coronae or coronae-like features (e.g., arachnoids), while the remainder are almost evenly distributed between the plains and highlands (e.g., Komatsu et al. 1993). Valley networks on Venus are also typically associated with coronae and sinuous rilles—and they seem to be subject to tectonic control than rilles (Komatsu et al. 2001). Overall, both rilles and valley networks appear to have formed via incision into the surrounding terrain. Other types of simple, complex, and compound channels appear shallower and often have flow margins—suggesting a constructional (i.e., volcanic) origin (e.g., Komatsu et al. 1993; Komatsu and Baker 1994; Komatsu et al. 2001).

Canali are long, narrow channels that superficially resemble meandering rivers on Earth. They are typically more than 500 km long, ~1 km wide, and ~24 m deep on average (Baker et al. 1992; Komatsu et al. 1993; Williams-Jones et al. 1998). Baltis Vallis is ~6,800 km long—the longest canali on Venus and the longest channel of any sort so far found in the solar system (e.g., Baker et al. 1992). This extreme length makes Baltis Vallis useful as a stratigraphic marker (e.g., Basilevsky and Head 1996) and as a tool with which to study the evolution of topography (e.g., Conrad and Nimmo 2023). River-like features such as cut-off meander bends, abandoned channel segments, and terminal deposits are often observed at canali. Narrow levees are sometimes seen, but only at the limit of the resolution of the radar imagery. Canali are concentrated on the plains—although identifying their sources and/or termini is often difficult or impossible. Unusual types of lava are often invoked to explain the apparently unique properties of canali (e.g., Kargel et al. 1994). For example, some studies favored carbonatite-rich (or carbonatite-sulphate-rich, with a eutectic composition) lavas that have low viscosity (to explain river-like meanders) and low melting temperatures (to explain the extreme lengths). For example, one carbonatite lava from the Oldoinyo Lengai volcano on Earth was observed to have a liquidus temperature of only 763 K, although ~1000 K is a more typical value of this type of lava (e.g., Dawson et al. 1990; Keller and Krafft 1990). Mafic and ultramafic lavas that are observed on Earth and other rocky planets, such as basalt or komatiite, have relatively high melting temperatures (>1300–1600 K). They could create channels via thermal erosion, but they might tend to solidify after travelling tens (or, at most, hundreds) of km across the surface (e.g., Flynn et al. 2023), rather than the thousands of km needed for some canali (e.g., Head and Wilson 1986).

Speculatively, the formation of canali may be important to models of the evolution of Venus. Carbonatite lavas on Earth typically contain >30–50 volume percent of carbonate minerals (Dawson 1962). Some studies suggest that canali formed with discharge rates >10⁵ m³ s⁻¹ for up to ~100 years (Kargel et al. 1994). Lava that formed each canale seen on the surface today thus could have degassed enough CO₂ to produce a small portion of the modern atmosphere (e.g., Kargel et al. 1994). Such rapid resurfacing with carbon-rich lava may lead to a coupled evolution of the surface and atmosphere—perhaps burying old impact craters (e.g., Komatsu et al. 1993; Strom et al. 1994) and creating a greenhouse effect that drives the global climate from clement to hellish (e.g., Way and Del Genio, 2020; Way et al., 2016). However, many questions about canali and carbonatite volcanism on Venus await answers.
plex channels without flow margins; and compound channels. Some of these channels resemble features seen on Mars and Earth's Moon. For example, sinuous rilles have large depth-to-width ratios and morphologies similar to those of lunar and Martian rilles (e.g., Komatsu and Baker 1994; Oshigami et al. 2009). Usually, no lateral flow deposits are observed near rilles, but flow deposits are sometimes found at their termini. Channel sources are sometimes unidentifiable, but they are often found to originate in depressions. Mappers have identified ~59 rilles with average lengths of ~10–300 km and widths up to several kilometres (e.g., Baker et al. 1992; Komatsu et al. 1993). More than half of the rilles are associated with coronae or coronae-like features (e.g., arachnoids), while the remainder are almost evenly distributed between the plains and highlands (e.g., Komatsu et al. 1993). Valley networks on Venus are also typically associated with coronae and sinuous rilles—and they seem to be subject to tectonic control than rilles (Komatsu et al. 2001). Overall, both rilles and valley networks appear to have formed via incision into the surrounding terrain. Other types of simple, complex, and compound channels appear shallower and often have flow margins—suggesting a constructional (i.e., volcanic) origin (e.g., Komatsu et al. 1993; Komatsu and Baker 1994; Komatsu et al. 2001).

Canali are long, narrow channels that superficially resemble meandering rivers on Earth. They are typically more than 500 km long, ~1 km wide, and ~24 m deep on average (Baker et al. 1992; Komatsu et al. 1993; Williams-Jones et al. 1998). Baltis Vallis is ~6,800 km long—the longest canali on Venus and the longest channel of any sort so far found in the solar system (e.g., Baker et al. 1992). This extreme length makes Baltis Vallis useful as a stratigraphic marker (e.g., Basilevsky and Head 1996) and as a tool with which to study the evolution of topography (e.g., Conrad and Nimmo 2023). River-like features such as cut-off meander bends, abandoned channel segments, and terminal deposits are often observed at canali. Narrow levees are sometimes seen, but only at the limit of the resolution of the radar imagery. Canali are concentrated on the plains—although identifying their sources and/or termini is often difficult or impossible. Unusual types of lava are often invoked to explain the apparently unique properties of canali (e.g., Kargel et al. 1994). For example, some studies favored carbonatite-rich (or carbonatite-sulphate-rich, with a eutectic composition) lavas that have low viscosity (to explain river-like meanders) and low melting temperatures (to explain the extreme lengths). For example, one carbonatite lava from the Oldoinyo Lengai volcano on Earth was observed to have a liquidus temperature of only 763 K, although ~1000 K is a more typical value of this type of lava (e.g., Dawson et al. 1990; Keller and Krafft 1990). Mafic and ultramafic lavas that are observed on Earth and other rocky planets, such as basalt or komatiite, have relatively high melting temperatures (>1300–1600 K). They could create channels via thermal erosion, but they might tend to solidify after travelling tens (or, at most, hundreds) of km across the surface (e.g., Flynn et al. 2023), rather than the thousands of km needed for some canali (e.g., Head and Wilson 1986).

Speculatively, the formation of canali may be important to models of the evolution of Venus. Carbonatite lavas on Earth typically contain >30–50 volume percent of carbonate minerals (Dawson 1962). Some studies suggest that canali formed with discharge rates >10⁵ m³ s⁻¹ for up to ~100 years (Kargel et al. 1994). Lava that formed each canale seen on the surface today thus could have degassed enough CO₂ to produce a small portion of the modern atmosphere (e.g., Kargel et al. 1994). Such rapid resurfacing with carbon-rich lava may lead to a coupled evolution of the surface and atmosphere—perhaps burying old impact craters (e.g., Komatsu et al. 1993; Strom et al. 1994) and creating a greenhouse effect that drives the global climate from clement to hellish (e.g., Way and Del Genio, 2020; Way et al., 2016). However, many questions about canali and carbonatite volcanism on Venus await answers.
## 2.1.2 Hotspots and Hotspot Volcanoes

Some of the large volcanoes on Venus can be classified as ‘hotspot’ volcanoes. On Earth, hotspots are defined as the surface manifestation of hot mantle that diapirically rises from the lower mantle or core–mantle boundary, generating volcanoes via pressure-release melting (Griffiths and Campbell 1990; Hoggard et al. 2020; Jellinek and Manga 2004; Morgan 1972). Mantle plumes and their tails are thought to be responsible for flood basalt volcanism and the formation of hotspots (White and McKenzie 1995); specifically, the bulbous head of a mantle plume impinges to the base of the lithosphere and generates mafic and ultramafic melts for a short time. On Earth, plumes have mantle potential temperatures that are 150 K to 350 K above ambient conditions. Seismic data indicates these mantle upwellings are generated at specific regions of the Earth’s lower mantle that are known as Large Low Shear Velocity Provinces (LLSVPs) (Torsvik et al., 2014).

Nine hotspots are inferred to have originated from or near the core–mantle boundary on Venus (McGill 1994; Smrekar et al. 1997). They have been identified on the basis of their great topographic swells (~1000–2000 km diameter, ~1–2 km high), large volcanic edifices, gravity anomalies that indicate compensation by low-density material at the base of the lithosphere (the plume head) (e.g., Smrekar and Phillips 1991), and, in some cases, rifts and coronae (Stofan et al. 1995). No hotspot tracks are observed in association with mantle plumes, consistent with a lack of plate tectonics operating at Venus. However, the large interior volcanoes and ring of wrinkle ridges surrounding Laufey Regio were interpreted as evidence of a past plume at that location (Brian et al. 2004). Once that rising mantle plume has cooled, the thermal isostasy from it no longer supports the topographic rise proposed to account for the circumferential wrinkle ridges.

Many coronae likely formed above small plumes but are typically too small to be caused by core–mantle boundary-sourced plumes. Artemis is larger (at ~2600 km in diameter) than many of the hotspot features discussed above, and has been proposed to be a site of plume-induced subduction (Davaille et al. 2017; McKenzie et al. 1992; Sandwell and Schubert 1992b). Other large coronae could have formed over deep mantle plumes: despite their relatively low resolution, gravity data at Quetzelpetlatl corona (~900 km diameter) are indicative of an active plume (Davaille et al. 2017). Heng-o (~1100 km diameter) is another large corona, but neither its topography nor gravity suggests activity today. The nine large volcanic rises (Stofan and Smrekar 2005) on Venus (ten, if Laufey is included) exclude these few very large coronae. This is based on both the smaller size of all features exclusive of Artemis and the analogy to Earth’s hotspots, which do not have the evidence of plume-induced subduction seen at Quetzelpetlatl and Artemis Coronae (Ivanov and Head, 2011; Davaille et al. 2017). However, plume-induced subduction is proposed to have occurred earlier in Earth’s history (e.g. Whattam and Stern, 2015), and thus excluding Artemis is perhaps artificial. Additionally, the larger features can be resolved in the Magellan gravity data, which provides evidence of an active plume at these features (e.g. Smrekar and Phillips 1991). As discussed below, there are multiple lines of evidence supporting these hotspots as likely active. Other (non-hotspot) volcanoes may also be active, but available data are inadequate to assess this possibility.

The classic volcanic edifices at terrestrial oceanic hotspots (e.g., Hawaii, Iceland) form shield volcanoes that are composed of a thick sequence (sometimes more than 10 km) of many individual lava flows (typically ~2–20 m thick). These flows in turn form a broad base, on the order of 100 km across, with relatively steep lower slopes (~10°) that grade to more gentle upper slopes (2° to 3°). The morphology of oceanic hotspot volcanoes is due to the sequential accumulation of low viscosity lava flows that erupt over a period of 1 million
## 2.1.2 Hotspots and Hotspot Volcanoes

Some of the large volcanoes on Venus can be classified as ‘hotspot’ volcanoes. On Earth, hotspots are defined as the surface manifestation of hot mantle that diapirically rises from the lower mantle or core–mantle boundary, generating volcanoes via pressure-release melting (Griffiths and Campbell 1990; Hoggard et al. 2020; Jellinek and Manga 2004; Morgan 1972). Mantle plumes and their tails are thought to be responsible for flood basalt volcanism and the formation of hotspots (White and McKenzie 1995); specifically, the bulbous head of a mantle plume impinges to the base of the lithosphere and generates mafic and ultramafic melts for a short time. On Earth, plumes have mantle potential temperatures that are 150 K to 350 K above ambient conditions. Seismic data indicates these mantle upwellings are generated at specific regions of the Earth’s lower mantle that are known as Large Low Shear Velocity Provinces (LLSVPs) (Torsvik et al., 2014).

Nine hotspots are inferred to have originated from or near the core–mantle boundary on Venus (McGill 1994; Smrekar et al. 1997). They have been identified on the basis of their great topographic swells (~1000–2000 km diameter, ~1–2 km high), large volcanic edifices, gravity anomalies that indicate compensation by low-density material at the base of the lithosphere (the plume head) (e.g., Smrekar and Phillips 1991), and, in some cases, rifts and coronae (Stofan et al. 1995). No hotspot tracks are observed in association with mantle plumes, consistent with a lack of plate tectonics operating at Venus. However, the large interior volcanoes and ring of wrinkle ridges surrounding Laufey Regio were interpreted as evidence of a past plume at that location (Brian et al. 2004). Once that rising mantle plume has cooled, the thermal isostasy from it no longer supports the topographic rise proposed to account for the circumferential wrinkle ridges.

Many coronae likely formed above small plumes but are typically too small to be caused by core–mantle boundary-sourced plumes. Artemis is larger (at ~2600 km in diameter) than many of the hotspot features discussed above, and has been proposed to be a site of plume-induced subduction (Davaille et al. 2017; McKenzie et al. 1992; Sandwell and Schubert 1992b). Other large coronae could have formed over deep mantle plumes: despite their relatively low resolution, gravity data at Quetzelpetlatl corona (~900 km diameter) are indicative of an active plume (Davaille et al. 2017). Heng-o (~1100 km diameter) is another large corona, but neither its topography nor gravity suggests activity today. The nine large volcanic rises (Stofan and Smrekar 2005) on Venus (ten, if Laufey is included) exclude these few very large coronae. This is based on both the smaller size of all features exclusive of Artemis and the analogy to Earth’s hotspots, which do not have the evidence of plume-induced subduction seen at Quetzelpetlatl and Artemis Coronae (Ivanov and Head, 2011; Davaille et al. 2017). However, plume-induced subduction is proposed to have occurred earlier in Earth’s history (e.g. Whattam and Stern, 2015), and thus excluding Artemis is perhaps artificial. Additionally, the larger features can be resolved in the Magellan gravity data, which provides evidence of an active plume at these features (e.g. Smrekar and Phillips 1991). As discussed below, there are multiple lines of evidence supporting these hotspots as likely active. Other (non-hotspot) volcanoes may also be active, but available data are inadequate to assess this possibility.

The classic volcanic edifices at terrestrial oceanic hotspots (e.g., Hawaii, Iceland) form shield volcanoes that are composed of a thick sequence (sometimes more than 10 km) of many individual lava flows (typically ~2–20 m thick). These flows in turn form a broad base, on the order of 100 km across, with relatively steep lower slopes (~10°) that grade to more gentle upper slopes (2° to 3°). The morphology of oceanic hotspot volcanoes is due to the sequential accumulation of low viscosity lava flows that erupt over a period of 1 million
years; volcanic activity can last up to 200 million years for a single hotspot (Heaman and Kjarsgaard 2000; Kasbohm and Schoene 2018; Wei et al. 2020). The lava compositions are 90 to 95% basaltic, with minor volumes of silicic volcanic rocks (e.g., rhyolite, trachyte, dacite) present within the youngest flow sequences (Jeffery and Gertisser 2018). After the initial effusive flows, the preponderance of subsequent flows during the shield-building stage are tholeiitic basalt, but minor volumes of alkali basalt may erupt during waning or rejuvenation stages (Thordarson and Garcia 2018).

On Venus, many of the hotspots described above have large shield volcanoes (e.g., Stofan et al. 1995), as well as other types of volcanism. Large shield volcanoes occur in many locations and geologic settings across Venus, as is the case on Earth (e.g., Hahn and Byrne, 2022). Without high-resolution geophysical and detailed geochemical data, the deep mantle origin and general eruption history of Venusian hotspot volcanism must be inferred from the geologic setting, gravity and topography, morphology and by analogy with Earth. Moreover, the quality of Magellan-based altimetric and stereo-derived topographic data are not sufficient to fully characterise volcano shape and slopes to the extent possible for Earth (and Mars). Nonetheless, on the basis of Magellan radar imagery, shield volcanoes on Venus appear to have broadly comparable morphologies and, presumably, slope values.

### 2.1.3 Coronae

Coronae are oblong to circular volcano-tectonic features ranging in diameter from 60 km to up to 2600 km, with a median size on the order of 200 km (Glaze et al. 2002). A ring of closely spaced concentric fractures and/or ridges is the defining characteristic of coronae. Stofan et al. (2001) divided coronae into two categories: Type 1, with >180° of a fracture annulus (Fig. 4c), and Type 2, or stealth coronae (Fig. 4b), with <180° (Stofan et al. 2001). Glaze et al. (2002) identified 406 coronae of Type 1 and 107 of Type 2. Coronae have also been categorized into five classes based on their fracture annuli: concentric, concentric double-ring, radial/concentric, asymmetric, and multiple (Stofan et al. 1992). Coronae often have circumferential rims or trenches that can extend beyond the annulus, which are not necessarily correlated radially or azimuthally with fracture annuli and can be equally incomplete and irregular in shape. They can also have interior domes, plateaus, and depressions. Based on these topographic elements, Smrekar and Stofan (1997) defined nine topographic classes. Three coronae that fall into a common topographic class (3a, rim surrounding interior high) with varying fracture patterns are illustrated in Fig. 4.

Coronae are found in every geologic environment on Venus (e.g., Stofan et al. 1992). The majority (62%) occur along fracture/rift belts (chasmata) (Glaze et al. 2002). Coronae appear randomly located in these extensional environments, suggesting that the detailed locations of coronae are not influenced by rifts and vice versa (Martin et al., 2007). Nonetheless, the concentration of coronae in extension environments implies a genetic relationship with these tectonic structures (e.g., Piskorz et al. 2014), although the remaining coronae occur in the plains (25%), at large topographic rises (11%), and tessera (1%) (Glaze et al. 2002).

Large volcanoes (see below) and smaller coronae overlap in scale. Morphologies can be transitional as well, leading some to propose that (some) coronae form via volcanic construction (Dombard et al. 2007; Lang and López 2015; McGovern et al. 2013). Others have pointed to the similarity between giant radiating dyke swarms on Earth and the radiating lineaments that occur at some coronae (e.g., Ernst et al. 1995).

The margins of some larger coronae, such as Artemis and Quetzalpetlal, have been proposed to be sites of retrograde lithospheric subduction (e.g., McKenzie et al. 1992; Sandwell
years; volcanic activity can last up to 200 million years for a single hotspot (Heaman and Kjarsgaard 2000; Kasbohm and Schoene 2018; Wei et al. 2020). The lava compositions are 90 to 95% basaltic, with minor volumes of silicic volcanic rocks (e.g., rhyolite, trachyte, dacite) present within the youngest flow sequences (Jeffery and Gertisser 2018). After the initial effusive flows, the preponderance of subsequent flows during the shield-building stage are tholeiitic basalt, but minor volumes of alkali basalt may erupt during waning or rejuvenation stages (Thordarson and Garcia 2018).

On Venus, many of the hotspots described above have large shield volcanoes (e.g., Stofan et al. 1995), as well as other types of volcanism. Large shield volcanoes occur in many locations and geologic settings across Venus, as is the case on Earth (e.g., Hahn and Byrne, 2022). Without high-resolution geophysical and detailed geochemical data, the deep mantle origin and general eruption history of Venusian hotspot volcanism must be inferred from the geologic setting, gravity and topography, morphology and by analogy with Earth. Moreover, the quality of Magellan-based altimetric and stereo-derived topographic data are not sufficient to fully characterise volcano shape and slopes to the extent possible for Earth (and Mars). Nonetheless, on the basis of Magellan radar imagery, shield volcanoes on Venus appear to have broadly comparable morphologies and, presumably, slope values.

### 2.1.3 Coronae

Coronae are oblong to circular volcano-tectonic features ranging in diameter from 60 km to up to 2600 km, with a median size on the order of 200 km (Glaze et al. 2002). A ring of closely spaced concentric fractures and/or ridges is the defining characteristic of coronae. Stofan et al. (2001) divided coronae into two categories: Type 1, with >180° of a fracture annulus (Fig. 4c), and Type 2, or stealth coronae (Fig. 4b), with <180° (Stofan et al. 2001). Glaze et al. (2002) identified 406 coronae of Type 1 and 107 of Type 2. Coronae have also been categorized into five classes based on their fracture annuli: concentric, concentric double-ring, radial/concentric, asymmetric, and multiple (Stofan et al. 1992). Coronae often have circumferential rims or trenches that can extend beyond the annulus, which are not necessarily correlated radially or azimuthally with fracture annuli and can be equally incomplete and irregular in shape. They can also have interior domes, plateaus, and depressions. Based on these topographic elements, Smrekar and Stofan (1997) defined nine topographic classes. Three coronae that fall into a common topographic class (3a, rim surrounding interior high) with varying fracture patterns are illustrated in Fig. 4.

Coronae are found in every geologic environment on Venus (e.g., Stofan et al. 1992). The majority (62%) occur along fracture/rift belts (chasmata) (Glaze et al. 2002). Coronae appear randomly located in these extensional environments, suggesting that the detailed locations of coronae are not influenced by rifts and vice versa (Martin et al., 2007). Nonetheless, the concentration of coronae in extension environments implies a genetic relationship with these tectonic structures (e.g., Piskorz et al. 2014), although the remaining coronae occur in the plains (25%), at large topographic rises (11%), and tessera (1%) (Glaze et al. 2002).

Large volcanoes (see below) and smaller coronae overlap in scale. Morphologies can be transitional as well, leading some to propose that (some) coronae form via volcanic construction (Dombard et al. 2007; Lang and López 2015; McGovern et al. 2013). Others have pointed to the similarity between giant radiating dyke swarms on Earth and the radiating lineaments that occur at some coronae (e.g., Ernst et al. 1995).

The margins of some larger coronae, such as Artemis and Quetzalpetlal, have been proposed to be sites of retrograde lithospheric subduction (e.g., McKenzie et al. 1992; Sandwell
<figure><img src="image_6.png" /></figure>

**Fig. 4** Magellan left-look SAR images of (a) Miralaidji, (b) Ponmakya, and (c) Aruru coronae, with respective stereo-derived topography (Herrick et al., 2012) in (d) and (e), and Magellan topography (f). Miralaidji Corona, near Latona and Dali Chasma, has radial fractures, and thus a morphology similar to that of novae (Krassilnikov and Head 2003). Ponmakya Corona, a stealth corona, has <180° concentric fractures, with a short NE-SW axis of ~175 km and long NW-SE axis of ~300 km, and >180° of a topographic rim. The appearance both in SAR and in topography is obscured by regional strain and interior volcanism, and orthogonal fractures obscure its appearance in SAR, in particular. Aruru Corona, a large volcanic centre (Tucker and Dombard 2023), has interior lava flooding, shield fields in the SW, and lava flows to the ESE

and Schubert 1992b; Schubert and Sandwell 1995). These coronae show topographic flexural signatures, beginning in the fracture annulus and extending outward, that resemble flexural signatures caused by retrograde subduction observed on Earth (Johnson and Sandwell 1994; O'Rourke and Smrekar 2018; Sandwell and Schubert 1992b). Moreover, interior extensional features displayed by Artemis corona are reminiscent of back-arc spreading centres (e.g., Brown and Grimm 1995; Sandwell and Schubert 1992b), contributing to the interpretation of retreating lithospheric subduction occurring at the south-eastern margin of the corona (McKenzie et al. 1992; Sandwell and Schubert 1992b), although others suggest Artemis Chasma is a site of convergence (Brown and Grimm 1995; Suppe and Connors 1992).

The diverse morphologies and sizes displayed by coronae may indicate that they form via a range of mechanisms. Modelling studies of the various mechanisms proposed to be responsible for corona formation are elaborated in more detail in Sect. 5.1.4 below. Regardless of the specific processes involved in their formation, coronae likely contribute to heat loss (Smrekar and Stofan 1997; Turcotte 1993). Moreover, coronae offer the opportunity to estimate the elastic lithospheric thickness through flexural modelling (e.g., O'Rourke and Smrekar 2018; Russell and Johnson 2021; Smrekar et al. 2023 and refer to Sect. 3 below and Rolf et al. 2022, this collection). As such, coronae offer important clues on the tectonic and resurfacing history of the planet.
<figure><img src="image_6.png" /></figure>

**Fig. 4** Magellan left-look SAR images of (a) Miralaidji, (b) Ponmakya, and (c) Aruru coronae, with respective stereo-derived topography (Herrick et al., 2012) in (d) and (e), and Magellan topography (f). Miralaidji Corona, near Latona and Dali Chasma, has radial fractures, and thus a morphology similar to that of novae (Krassilnikov and Head 2003). Ponmakya Corona, a stealth corona, has <180° concentric fractures, with a short NE-SW axis of ~175 km and long NW-SE axis of ~300 km, and >180° of a topographic rim. The appearance both in SAR and in topography is obscured by regional strain and interior volcanism, and orthogonal fractures obscure its appearance in SAR, in particular. Aruru Corona, a large volcanic centre (Tucker and Dombard 2023), has interior lava flooding, shield fields in the SW, and lava flows to the ESE.

and Schubert 1992b; Schubert and Sandwell 1995). These coronae show topographic flexural signatures, beginning in the fracture annulus and extending outward, that resemble flexural signatures caused by retrograde subduction observed on Earth (Johnson and Sandwell 1994; O'Rourke and Smrekar 2018; Sandwell and Schubert 1992b). Moreover, interior extensional features displayed by Artemis corona are reminiscent of back-arc spreading centres (e.g., Brown and Grimm 1995; Sandwell and Schubert 1992b), contributing to the interpretation of retreating lithospheric subduction occurring at the south-eastern margin of the corona (McKenzie et al. 1992; Sandwell and Schubert 1992b), although others suggest Artemis Chasma is a site of convergence (Brown and Grimm 1995; Suppe and Connors 1992).

The diverse morphologies and sizes displayed by coronae may indicate that they form via a range of mechanisms. Modelling studies of the various mechanisms proposed to be responsible for corona formation are elaborated in more detail in Sect. 5.1.4 below. Regardless of the specific processes involved in their formation, coronae likely contribute to heat loss (Smrekar and Stofan 1997; Turcotte 1993). Moreover, coronae offer the opportunity to estimate the elastic lithospheric thickness through flexural modelling (e.g., O'Rourke and Smrekar 2018; Russell and Johnson 2021; Smrekar et al. 2023 and refer to Sect. 3 below and Rolf et al. 2022, this collection). As such, coronae offer important clues on the tectonic and resurfacing history of the planet.
<figure><img src="image_7.png" /><figcaption>**Fig. 5** - Magellan topography data superposed by locations of topographic rises (labelled with white names), major crustal plateaus (labelled with black names), Type 1 coronae (red dots), and Type 2 coronae (white dots). Blues and purple are topographically low regions; yellow and white are topographically high. Coronae locations and classifications from Stofan et al. (2001); Glaze et al. (2002)</figcaption></figure>

### 2.1.4 Novae

Novae (singular: nova) are structures with prominent, stellate fracture patterns consisting of a mesh of graben centered on a central summit (Head et al. 1992; Krassilnikov and Head 2003). They are usually 100–300 km in diameter, have a prominent upraised topography (refer to Fig. 4a), and are associated with lava flows. Apparently distinct from volcanoes, novae show a dominance of tectonic over volcanic activity. Sixty-four novae have been identified on the Venusian surface (Krassilnikov and Head 2003), mainly located in areas of regional rises and rift zones, with a small number in the lowlands.

It is generally assumed that novae are formed by volcanic updoming and fracturing due to impingement of hot mantle diapirs on the upper part of the lithosphere (Janes et al. 1992; Koch and Manga 1996). Novae have been proposed to be the initial stages of corona development (e.g., Janes et al. 1992), or even perhaps “failed coronae” if the source of dynamic support were retracted early in the formation of what would go on to become a corona. Some studies suggested that novae within coronae represent reactivation of the corona location (Aittola and Kostama 2002), while others pointed out that the lengths of the radial structures often exceed the anticipated distance for diapiric uplift and fracturing, and instead linked novae to radial fracturing caused by dyke emplacement processes from a central magma source (e.g., Parfitt and Head 1993). However, it is important to note that <5% of coronae have radial fractures similar to novae, calling into question the hypothesis that novae are the first stage of corona formation.

### 2.2 Evidence for Recent Volcanism and Likely Active Features

One of the biggest scientific debates about Venus concerns its style of volcanism throughout its geologic history. As discussed above and in Herrick et al. (2023, this collection),
<figure><img src="image_7.png" /><figcaption>**Fig. 5** - Magellan topography data superposed by locations of topographic rises (labelled with white names), major crustal plateaus (labelled with black names), Type 1 coronae (red dots), and Type 2 coronae (white dots). Blues and purple are topographically low regions; yellow and white are topographically high. Coronae locations and classifications from Stofan et al. (2001); Glaze et al. (2002)</figcaption></figure>

### 2.1.4 Novae

Novae (singular: nova) are structures with prominent, stellate fracture patterns consisting of a mesh of graben centered on a central summit (Head et al. 1992; Krassilnikov and Head 2003). They are usually 100–300 km in diameter, have a prominent upraised topography (refer to Fig. 4a), and are associated with lava flows. Apparently distinct from volcanoes, novae show a dominance of tectonic over volcanic activity. Sixty-four novae have been identified on the Venusian surface (Krassilnikov and Head 2003), mainly located in areas of regional rises and rift zones, with a small number in the lowlands.

It is generally assumed that novae are formed by volcanic updoming and fracturing due to impingement of hot mantle diapirs on the upper part of the lithosphere (Janes et al. 1992; Koch and Manga 1996). Novae have been proposed to be the initial stages of corona development (e.g., Janes et al. 1992), or even perhaps “failed coronae” if the source of dynamic support were retracted early in the formation of what would go on to become a corona. Some studies suggested that novae within coronae represent reactivation of the corona location (Aittola and Kostama 2002), while others pointed out that the lengths of the radial structures often exceed the anticipated distance for diapiric uplift and fracturing, and instead linked novae to radial fracturing caused by dyke emplacement processes from a central magma source (e.g., Parfitt and Head 1993). However, it is important to note that <5% of coronae have radial fractures similar to novae, calling into question the hypothesis that novae are the first stage of corona formation.

## 2.2 Evidence for Recent Volcanism and Likely Active Features

One of the biggest scientific debates about Venus concerns its style of volcanism throughout its geologic history. As discussed above and in Herrick et al. (2023, this collection),
both catastrophic and regional resurfacing models are viable for Venus. Rolf et al. (2022, this collection) discuss the importance of understanding the history of volcanism for internal evolution models. Over the last ~15 years, numerous studies have shown evidence for current volcanic activity, suggesting Venus is likely more active today than previously understood. Many studies have argued for recent volcanism based on signatures of chemical, thermal, atmospheric gas, and surface change. Analysis of geophysical data supports the location of activity at rifts and hotspots.

Chemical signatures of recent volcanism are based on interpreting surface emissivity data as an indication that the weathering of new lava flows has not proceeded to completion. Evidence of 'recent' volcanism was first noted in Magellan emissivity data. The relative youth of Idunn and Maat Montes as well as some coronae is further corroborated by high values of Magellan radar emissivity at their summits, consistent with relatively little weathering (Klose et al. 1992; Robinson and Wood 1993). Brossier et al. (2020) further investigated these and other regions, suggesting additional minerals that could explain these signatures. Measurements made by the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) instrument aboard Venus Express were used to derive a map of thermal emission at 1.02 microns of much of Venus' southern hemisphere added further evidence of recent volcanism. Higher thermal emissivity displayed by some lava flows were interpreted as the chemical weathering signature of geologically recent lava flows (Smrekar et al. 2010; Stofan et al. 2016; D'Incecco et al. 2017). The presence of volcanic activity was found at Idunn Mons (in Imdr Regio), Hathor and Innini Montes (in Dione Regio), and Mielikki Mons (in Themis Regio). Prior interpretation of gravity and topography data in these areas provided evidence of mantle plumes beneath these regions, providing further evidence of activity in these areas (Smrekar et al. 2010 and references therein). The age of these incompletely weathered flows is debated. Filiberto et al. (2020) suggest that the flows are as young as a few years based on weathering experiments done using terrestrial air. Others argue for much slower weathering (up to ~500,000 yr) based on kinetic calculations applied to Venus spectral data (Dyar et al. 2020). Much work remains to fully interpret the age of 'recent' flows.

Other studies have provided evidence of a thermal signature of active flows. Transient bright spots in data from the Venus Monitoring Camera (VMC) were suggested to arise from volcanic activity associated with the Ganis Chasma rift zone in the Atla Regio (Shalygin et al. 2015). Some transient features could also be artefacts of cloud motion or instrumental corrections for background temperature. However, the areas with the bright spots also have high radar emissivity associated with relative geologic youthfulness (Brossier et al. 2020). Bondarenko et al. (2010) argued for subsurface thermal anomalies in Bereghinia Planitia using Magellan microwave radiometer data. This area, along with ~50% of the planet, is not covered by VIRTIS data, which cover most of the southern hemisphere. Additionally, the Magellan radar emissivity signatures, which may be due to incomplete weathering, are found only above ~0.7 km altitude.

Together with Earth-based radar maps, Campbell et al. (2017) interpreted several radar-bright deposits to reflect the early stage of recently renewed magmatic activity in eastern Eistla, western Eistla, Phoebe, and Dione Regiones. Herrick and Hensley (2023) found a volcanic vent near the summit of Maat Mons that seems to have changed in shape during the Magellan mission, providing the first direct observational evidence for active surface change driven by volcanism. Both studies offer evidence for volcanism using radar imaging in the same areas where Magellan radar and/or VIRTIS thermal emissivity offer evidence of incomplete weathering.

Interpretation of the gravity and topography data (See Sect. 4) is also consistent with the locations of many of these recent and perhaps currently active centers of volcanism
both catastrophic and regional resurfacing models are viable for Venus. Rolf et al. (2022, this collection) discuss the importance of understanding the history of volcanism for internal evolution models. Over the last ~15 years, numerous studies have shown evidence for current volcanic activity, suggesting Venus is likely more active today than previously understood. Many studies have argued for recent volcanism based on signatures of chemical, thermal, atmospheric gas, and surface change. Analysis of geophysical data supports the location of activity at rifts and hotspots.

Chemical signatures of recent volcanism are based on interpreting surface emissivity data as an indication that the weathering of new lava flows has not proceeded to completion. Evidence of 'recent' volcanism was first noted in Magellan emissivity data. The relative youth of Idunn and Maat Montes as well as some coronae is further corroborated by high values of Magellan radar emissivity at their summits, consistent with relatively little weathering (Klose et al. 1992; Robinson and Wood 1993). Brossier et al. (2020) further investigated these and other regions, suggesting additional minerals that could explain these signatures. Measurements made by the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) instrument aboard Venus Express were used to derive a map of thermal emission at 1.02 microns of much of Venus' southern hemisphere added further evidence of recent volcanism. Higher thermal emissivity displayed by some lava flows were interpreted as the chemical weathering signature of geologically recent lava flows (Smrekar et al. 2010; Stofan et al. 2016; D'Incecco et al. 2017). The presence of volcanic activity was found at Idunn Mons (in Imdr Regio), Hathor and Innini Montes (in Dione Regio), and Mielikki Mons (in Themis Regio). Prior interpretation of gravity and topography data in these areas provided evidence of mantle plumes beneath these regions, providing further evidence of activity in these areas (Smrekar et al. 2010 and references therein). The age of these incompletely weathered flows is debated. Filiberto et al. (2020) suggest that the flows are as young as a few years based on weathering experiments done using terrestrial air. Others argue for much slower weathering (up to ~500,000 yr) based on kinetic calculations applied to Venus spectral data (Dyar et al. 2020). Much work remains to fully interpret the age of 'recent' flows.

Other studies have provided evidence of a thermal signature of active flows. Transient bright spots in data from the Venus Monitoring Camera (VMC) were suggested to arise from volcanic activity associated with the Ganis Chasma rift zone in the Atla Regio (Shalygin et al. 2015). Some transient features could also be artefacts of cloud motion or instrumental corrections for background temperature. However, the areas with the bright spots also have high radar emissivity associated with relative geologic youthfulness (Brossier et al. 2020). Bondarenko et al. (2010) argued for subsurface thermal anomalies in Bereghinia Planitia using Magellan microwave radiometer data. This area, along with ~50% of the planet, is not covered by VIRTIS data, which cover most of the southern hemisphere. Additionally, the Magellan radar emissivity signatures, which may be due to incomplete weathering, are found only above ~0.7 km altitude.

Together with Earth-based radar maps, Campbell et al. (2017) interpreted several radar-bright deposits to reflect the early stage of recently renewed magmatic activity in eastern Eistla, western Eistla, Phoebe, and Dione Regiones. Herrick and Hensley (2023) found a volcanic vent near the summit of Maat Mons that seems to have changed in shape during the Magellan mission, providing the first direct observational evidence for active surface change driven by volcanism. Both studies offer evidence for volcanism using radar imaging in the same areas where Magellan radar and/or VIRTIS thermal emissivity offer evidence of incomplete weathering.

Interpretation of the gravity and topography data (See Sect. 4) is also consistent with the locations of many of these recent and perhaps currently active centers of volcanism
by revealing regions of upwelling and thin lithosphere. The long-wavelength geoid data suggest that both rifts and many areas that exhibit wrinkle ridges are likely supported by dynamic compensation and thermal anomalies at depth (Sandwell et al., 1997). The strong spatial correlation of the transient bright spots with rift zones suggests a correlation between deep and broad thermal upwellings and surface volcanism (Piskorz et al. 2014; Shalygin et al. 2015). Smrekar et al. (2010) provide an overview of several studies pointing to mantle plumes beneath the volcanic centers where emissivity anomalies are located. Other studies estimate elastic thickness from topographic flexure or admittance, and then derive heat flow as discussed in Sect. 3.2. High heat flow is a strong indication of activity, and is found at most coronae and nearby features where flexure is observed, including those coronae concentrated along Parga Chasma (e.g., O'Rourke and Smrekar 2018, Russell and Johnson 2021, Smrekar et al. 2023). Additionally, Gülcher et al. (2020) argued that 37 coronae with outer trench and rise morphology require ongoing suction above downwards-moving lithospheric material and an elevated interior supported by upwelling buoyancy.

Atmospheric data also offer a whiff of recent volcanism at unspecified sites. Varying rates of atmospheric SO₂ were interpreted as evidence of volcanic activity (e.g., Marcq et al. 2013). Analysis of the Soviet VeGa-2 probe (1985) data indicated, in the last few kilometres above the ground level, a lapse rate that would be unstable in the absence of a stabilis- ing chemical gradient, which has been suggested to arise from a vertical gradient in nitro- gen caused by density-driven separation of N₂ from CO₂ (Lebonnois and Schubert 2017). Cordier et al. (2019) found that producing a stabilising gradient with only CO₂ released from the Venusian crust would require an unrealistically huge rate of volcanic degassing. See also Wilson et al. (2024, this collection) for a discussion of effects of volcanic eruptions on the modern atmosphere of Venus.

# 3 Overview of Tectonic Features

## 3.1 Tessera Terrain

Approximately 8% of the surface of Venus is classified as tessera terrain (Ivanov and Head 1996), defined as having relatively complex patterns of deformation (e.g., Barsukov et al. 1986; Ivanov and Head; 1996). Tesserae are found as large (~1000 s km) plateaus stand- ing ~1–4 km above the mean planetary radius, and as small outcrops (~10 s –100 s km) scattered in the plains (Gilmore et al., 2023, this collection). Global mapping of the tessera shows that most tessera margins are embayed by plains, thus establishing the tessera as the oldest materials in a proposed global stratigraphic column (e.g., Ivanov and Head 2011, 2013). This interpretation is supported by the total crater population of tessera terrain, which with ~80 craters is up to ~1.4 times the global average crater age (Gilmore et al. 1997; Ivanov and Basilevsky 1993), and may be older if deformed craters are recognized with fu- ture, higher-resolution imaging and topography data. The crater record, while sparse, does indicate that the observable phases of tessera deformation occurred at the time of, or just before, the global average age, although the rock age of the tesserae is unknown. Some tessera boundaries show interaction with plains units. Margin-parallel tessera structures are observed to uplift and deformed later plains units, e.g., in the southeast part of Thetis Re- gio (Ghail 2002), the western boundary of Alpha Regio (Gilmore and Head 2000), and the northern margin of Ovda Regio (Romeo and Capote 2011), suggesting that tessera deforma- tion continued in places after the emplacement of marginal plains.
by revealing regions of upwelling and thin lithosphere. The long-wavelength geoid data suggest that both rifts and many areas that exhibit wrinkle ridges are likely supported by dynamic compensation and thermal anomalies at depth (Sandwell et al., 1997). The strong spatial correlation of the transient bright spots with rift zones suggests a correlation between deep and broad thermal upwellings and surface volcanism (Piskorz et al. 2014; Shalygin et al. 2015). Smrekar et al. (2010) provide an overview of several studies pointing to mantle plumes beneath the volcanic centers where emissivity anomalies are located. Other studies estimate elastic thickness from topographic flexure or admittance, and then derive heat flow as discussed in Sect. 3.2. High heat flow is a strong indication of activity, and is found at most coronae and nearby features where flexure is observed, including those coronae concentrated along Parga Chasma (e.g., O'Rourke and Smrekar 2018, Russell and Johnson 2021, Smrekar et al. 2023). Additionally, Gülcher et al. (2020) argued that 37 coronae with outer trench and rise morphology require ongoing suction above downwards-moving lithospheric material and an elevated interior supported by upwelling buoyancy.

Atmospheric data also offer a whiff of recent volcanism at unspecified sites. Varying rates of atmospheric SO₂ were interpreted as evidence of volcanic activity (e.g., Marcq et al. 2013). Analysis of the Soviet VeGa-2 probe (1985) data indicated, in the last few kilometres above the ground level, a lapse rate that would be unstable in the absence of a stabilis- ing chemical gradient, which has been suggested to arise from a vertical gradient in nitro- gen caused by density-driven separation of N₂ from CO₂ (Lebonnois and Schubert 2017). Cordier et al. (2019) found that producing a stabilising gradient with only CO₂ released from the Venusian crust would require an unrealistically huge rate of volcanic degassing. See also Wilson et al. (2024, this collection) for a discussion of effects of volcanic eruptions on the modern atmosphere of Venus.

# 3 Overview of Tectonic Features

## 3.1 Tessera Terrain

Approximately 8% of the surface of Venus is classified as tessera terrain (Ivanov and Head 1996), defined as having relatively complex patterns of deformation (e.g., Barsukov et al. 1986; Ivanov and Head; 1996). Tesserae are found as large (~1000 s km) plateaus stand- ing ~1–4 km above the mean planetary radius, and as small outcrops (~10 s –100 s km) scattered in the plains (Gilmore et al., 2023, this collection). Global mapping of the tessera shows that most tessera margins are embayed by plains, thus establishing the tessera as the oldest materials in a proposed global stratigraphic column (e.g., Ivanov and Head 2011, 2013). This interpretation is supported by the total crater population of tessera terrain, which with ~80 craters is up to ~1.4 times the global average crater age (Gilmore et al. 1997; Ivanov and Basilevsky 1993), and may be older if deformed craters are recognized with fu- ture, higher-resolution imaging and topography data. The crater record, while sparse, does indicate that the observable phases of tessera deformation occurred at the time of, or just before, the global average age, although the rock age of the tesserae is unknown. Some tessera boundaries show interaction with plains units. Margin-parallel tessera structures are observed to uplift and deformed later plains units, e.g., in the southeast part of Thetis Re- gio (Ghail 2002), the western boundary of Alpha Regio (Gilmore and Head 2000), and the northern margin of Ovda Regio (Romeo and Capote 2011), suggesting that tessera deforma- tion continued in places after the emplacement of marginal plains.
<figure><img src="image_8.png" /><figcaption>**Fig. 6** - Margin-parallel ridge belts adjacent to interior facies of Ovda Regio</figcaption></figure>

Shortening, extensional and transpressional structures are recognized in the tesserae (e.g., Ghent and Hansen 1999; Gilmore et al. 1998; Tuckwell and Ghail 2003). The style of extensional structures affects the interpretation of plateau origin and stratigraphy (i.e., whether their formation started with tectonic compression versus extension). First, these extensional features may be graben resulting from crustal relaxation that deformed older compression structures that, in turn, originated presumably from an earlier compressional, plateau-building phase (e.g., Bindschadler et al. 1992). Second, an alternative interpretation holds that the extensional features are steep (~90°), open tensile fractures (called ribbon terrain). This hypothesis requires that they formed during an early extensional phase that precedes later compression and wide graben formation (e.g., Ghent and Tibuleac 2002; Hansen and Willis 1996). These endmember models call for the formation of tessera plateaus during mantle downwelling or upwelling (e.g., Phillips and Hansen 1998; Gilmore et al. 1998; Hansen et al. 2000), respectively, which can be tested with high-resolution images of extensional features. In several large plateaus (e.g., Ovda Regio, Tellus Regio), ridge belts are found parallel to the plateau margins, surrounding an interior with a different structural fabric (Fig. 6). At Tellus Regio, specifically, lateral accretion of the margins possibly occurred after the formation of the structures in the facies of the plateau interior (Gilmore and Head
<figure><img src="image_8.png" /><figcaption>**Fig. 6** - Margin-parallel ridge belts adjacent to interior facies of Ovda Regio</figcaption></figure>

Shortening, extensional and transpressional structures are recognized in the tesserae (e.g., Ghent and Hansen 1999; Gilmore et al. 1998; Tuckwell and Ghail 2003). The style of extensional structures affects the interpretation of plateau origin and stratigraphy (i.e., whether their formation started with tectonic compression versus extension). First, these extensional features may be graben resulting from crustal relaxation that deformed older compression structures that, in turn, originated presumably from an earlier compressional, plateau-building phase (e.g., Bindschadler et al. 1992). Second, an alternative interpretation holds that the extensional features are steep (~90°), open tensile fractures (called ribbon terrain). This hypothesis requires that they formed during an early extensional phase that precedes later compression and wide graben formation (e.g., Ghent and Tibuleac 2002; Hansen and Willis 1996). These endmember models call for the formation of tessera plateaus during mantle downwelling or upwelling (e.g., Phillips and Hansen 1998; Gilmore et al. 1998; Hansen et al. 2000), respectively, which can be tested with high-resolution images of extensional features. In several large plateaus (e.g., Ovda Regio, Tellus Regio), ridge belts are found parallel to the plateau margins, surrounding an interior with a different structural fabric (Fig. 6). At Tellus Regio, specifically, lateral accretion of the margins possibly occurred after the formation of the structures in the facies of the plateau interior (Gilmore and Head
2018; Resor et al. 2021). Therefore, at least some tessera facies possibly formed prior to their incorporation via lateral accretion—and subsequent deformation during plateau formation.

Several lines of evidence show that tessera formation is linked to an extinct geodynamic regime. Buckle analysis of the wavelength of tesserae folds show that they require both elevated heat flows and strain rates relative to structures preserved in the plains, coronae, and volcanoes (e.g., Brown and Grimm, 1997; 1999). This interpretation is valid for dry rocks of basaltic through felsic compositions if the rocks are not quartz-dominated (Resor et al. 2021). Multiple wavelengths of deformation might alternatively reflect physical layering in the crust (e.g., Romeo and Capote 2011), although this explanation may require layers that are thicker than observed (e.g., Byrne et al. 2021). The tesserae (average crustal thickness of ~20 km) are also currently isostatically compensated, again consistent with a lack of recent mantle contribution (e.g., Anderson and Smrekar, 2006; Maia and Wieczorek, 2022).

The composition of the tesserae has not yet been measured directly, but that of some tessera exposures can be inferred from both NIR and radar observations. Observations at ~1–1.2 µm from the Galileo flyby (Hashimoto et al. 2008) and VEX VIRTIS (Mueller et al. 2008; Gilmore et al. 2015) show that the Alpha Regio tessera has a lower emissivity than the global average, which is dominated by plains of presumably basaltic composition. Since emissivity near 1 µm is inversely correlated to Fe²⁺ content in minerals (e.g., Hunt and Salisbury 1970; Dyar et al. 2020), the signature of the tessera is consistent with rocks with lower Fe content, or more felsic compositions, assuming they are igneous rocks (e.g., Gilmore et al. 2023). The generation of substantial amounts of felsic melt (i.e., on the order of the volume of the tesserae) could require both water and a lithospheric recycling mechanism for formation (e.g., Campbell and Taylor 1983), and thus possibly must have formed during a more habitable era if they are ultimately shown to be chemically equivalent to Earth’s continental crust (e.g., Campbell and Taylor 1983; Gilmore et al. 2023).

The interpretations above generally assume that the tesserae are igneous in origin. The complex structural fabric of tesserae makes it difficult to document the morphology of the original surfaces at the scale of the Magellan imagery, although Ivanov (2001) asserted from a survey that tessera precursor terrain is smooth and plains-like. However, layered beds have been identified in Tellus Regio, which resemble sedimentary beds on Earth and Mars exposed by erosion (Byrne et al., 2021). Additionally, Khawja et al. (2020) proposed that the topography of some tesserae is similar to fluvial drainage basins. It also should be recognized that the tesserae need not be structurally, morphologically, compositionally, or stratigraphically homogeneous: e.g., Brossier et al. (2020) showed that the radar emissivity properties of tesserae vary geographically, which may be evidence for differences in rock composition. The tesserae are also likely to contain sedimentary deposits from nearby, upwind craters (e.g., Campbell et al. 2015; Whitten and Campbell 2016) and internal mass wasting (e.g., Carter et al. 2023, this collection). The imaging, topography, and gravity data provided from VERITAS, DAVINCI and EnVision (Widemann et al. 2023, this collection; see also Sect. 6) will substantially advance our understanding of this important terrain type.

## 3.2 Rifts, Extension and Dykes

Long systems of extensional structures on Venus have variously been termed “fracture belts” or “groove belts” in the literature (e.g., Squyres et al. 1992; Ivanov and Head 2011) but, by analogy with Earth, are sites of crustal extension or rifting. These systems host normal faults that form graben and half graben with abundant evidence for fault linkage. Many of these systems are tens of kilometres across and hundreds of kilometres long. At larger scales, belts of extensional structures up to ten thousand kilometres long (Fig. 6) have been
2018; Resor et al. 2021). Therefore, at least some tessera facies possibly formed prior to their incorporation via lateral accretion—and subsequent deformation during plateau formation.

Several lines of evidence show that tessera formation is linked to an extinct geodynamic regime. Buckle analysis of the wavelength of tessera folds show that they require both elevated heat flows and strain rates relative to structures preserved in the plains, coronae, and volcanoes (e.g., Brown and Grimm, 1997; 1999). This interpretation is valid for dry rocks of basaltic through felsic compositions if the rocks are not quartz-dominated (Resor et al. 2021). Multiple wavelengths of deformation might alternatively reflect physical layering in the crust (e.g., Romeo and Capote 2011), although this explanation may require layers that are thicker than observed (e.g., Byrne et al. 2021). The tesserae (average crustal thickness of ~20 km) are also currently isostatically compensated, again consistent with a lack of recent mantle contribution (e.g., Anderson and Smrekar, 2006; Maia and Wieczorek, 2022).

The composition of the tesserae has not yet been measured directly, but that of some tessera exposures can be inferred from both NIR and radar observations. Observations at ~1–1.2 µm from the Galileo flyby (Hashimoto et al. 2008) and VEX VIRTIS (Mueller et al. 2008; Gilmore et al. 2015) show that the Alpha Regio tessera has a lower emissivity than the global average, which is dominated by plains of presumably basaltic composition. Since emissivity near 1 µm is inversely correlated to Fe²⁺ content in minerals (e.g., Hunt and Salisbury 1970; Dyar et al. 2020), the signature of the tessera is consistent with rocks with lower Fe content, or more felsic compositions, assuming they are igneous rocks (e.g., Gilmore et al. 2023). The generation of substantial amounts of felsic melt (i.e., on the order of the volume of the tesserae) could require both water and a lithospheric recycling mechanism for formation (e.g., Campbell and Taylor 1983), and thus possibly must have formed during a more habitable era if they are ultimately shown to be chemically equivalent to Earth’s continental crust (e.g., Campbell and Taylor 1983; Gilmore et al. 2023).

The interpretations above generally assume that the tesserae are igneous in origin. The complex structural fabric of tesserae makes it difficult to document the morphology of the original surfaces at the scale of the Magellan imagery, although Ivanov (2001) asserted from a survey that tessera precursor terrain is smooth and plains-like. However, layered beds have been identified in Tellus Regio, which resemble sedimentary beds on Earth and Mars exposed by erosion (Byrne et al., 2021). Additionally, Khawja et al. (2020) proposed that the topography of some tesserae is similar to fluvial drainage basins. It also should be recognized that the tesserae need not be structurally, morphologically, compositionally, or stratigraphically homogeneous: e.g., Brossier et al. (2020) showed that the radar emissivity properties of tesserae vary geographically, which may be evidence for differences in rock composition. The tesserae are also likely to contain sedimentary deposits from nearby, upwind craters (e.g., Campbell et al. 2015; Whitten and Campbell 2016) and internal mass wasting (e.g., Carter et al. 2023, this collection). The imaging, topography, and gravity data provided from VERITAS, DAVINCI and EnVision (Widemann et al. 2023, this collection; see also Sect. 6) will substantially advance our understanding of this important terrain type.

## 3.2 Rifts, Extension and Dykes

Long systems of extensional structures on Venus have variously been termed “fracture belts” or “groove belts” in the literature (e.g., Squyres et al. 1992; Ivanov and Head 2011) but, by analogy with Earth, are sites of crustal extension or rifting. These systems host normal faults that form graben and half graben with abundant evidence for fault linkage. Many of these systems are tens of kilometres across and hundreds of kilometres long. At larger scales, belts of extensional structures up to ten thousand kilometres long (Fig. 6) have been
<figure><img src="image_9.png" /><figcaption>**Fig. 7** - Parga Chasma, trending from NW to SE, or Atla Regio to Themis Regio, appears bright in Magellan SAR imagery. Note dozens of coronae in proximity to the rift</figcaption></figure>

assigned the term chasmata or “rift zone” based on analogies to features on Earth and Mars (e.g., Solomon et al. 1991, 1992; Ivanov and Head 2011). These rifts form deep, elongated troughs, generally accompanied by broad expanses of effusive volcanic deposits, but are themselves typically located within regional topographic highs.

A large fraction of coronae are situated in proximity to the major rifts of Parga (Fig. 7), Hecate, and Dali/Diana (Hamilton and Stofan, 1996; Martin et al., 2007); many more occur in association with smaller rift zones (Glaze et al. 2002) (Sect. 2.1.3). Coronae at large rifts were found to have various age relationships with the rifts, such as coronae younger than the rift, synchronous formation with the rift, and post-tectonic (Hamilton and Stofan, 1996; Martin et al. 2007). At least one rift zone, Ganis Chasma, may be the site of active volcanism today (Shalygin et al. 2015). The collocation of rifts on Venus with elevated topography, together with considerable depths of compensation given by gravity-topography admittance functions, suggests that they are supported at least in part by mantle upwelling (e.g., Solomon et al. 1991; Smrekar et al. 2010). There are numerous examples of radiating extensional fracture systems on Venus (Grosfils and Head, 1994b) (Sect. 2.1.4), with at least 100 morphologically consistent with being associated with, and perhaps formed by, dyke emplacement (Ernst et al., 2001). Several of these (putative) radiating dyke swarms on Venus are associated with broad, domical swells comparable in size to mantle upwelling-induced uplifts on Earth (Grosfils and Head, 1994b).

## 3.3 Ridge Belts and Wrinkle Ridges

Where crustal shortening is spatially concentrated on Venus, such deformation is typically manifest as bands of structures interpreted as thrust faults and folds, and which have been termed “ridge belts” in the literature (e.g., Barsukov et al. 1986; Squyres et al., 1992, see Fig. 8). This term is not an established geologic one. Although it may be useful in certain circumstances, using more general terms such as “mountain ranges” may better align the study of Venus geology with the other terrestrial worlds (Klimczak et al., 2019). In any
<figure><img src="image_9.png" /><figcaption>**Fig. 7** - Parga Chasma, trending from NW to SE, or Atla Regio to Themis Regio, appears bright in Magellan SAR imagery. Note dozens of coronae in proximity to the rift</figcaption></figure>

assigned the term chasmata or “rift zone” based on analogies to features on Earth and Mars (e.g., Solomon et al. 1991, 1992; Ivanov and Head 2011). These rifts form deep, elongated troughs, generally accompanied by broad expanses of effusive volcanic deposits, but are themselves typically located within regional topographic highs.

A large fraction of coronae are situated in proximity to the major rifts of Parga (Fig. 7), Hecate, and Dali/Diana (Hamilton and Stofan, 1996; Martin et al., 2007); many more occur in association with smaller rift zones (Glaze et al. 2002) (Sect. 2.1.3). Coronae at large rifts were found to have various age relationships with the rifts, such as coronae younger than the rift, synchronous formation with the rift, and post-tectonic (Hamilton and Stofan, 1996; Martin et al. 2007). At least one rift zone, Ganis Chasma, may be the site of active volcanism today (Shalygin et al. 2015). The collocation of rifts on Venus with elevated topography, together with considerable depths of compensation given by gravity-topography admittance functions, suggests that they are supported at least in part by mantle upwelling (e.g., Solomon et al. 1991; Smrekar et al. 2010). There are numerous examples of radiating extensional fracture systems on Venus (Grosfils and Head, 1994b) (Sect. 2.1.4), with at least 100 morphologically consistent with being associated with, and perhaps formed by, dyke emplacement (Ernst et al., 2001). Several of these (putative) radiating dyke swarms on Venus are associated with broad, domical swells comparable in size to mantle upwelling-induced uplifts on Earth (Grosfils and Head, 1994b).

### 3.3 Ridge Belts and Wrinkle Ridges

Where crustal shortening is spatially concentrated on Venus, such deformation is typically manifest as bands of structures interpreted as thrust faults and folds, and which have been termed “ridge belts” in the literature (e.g., Barsukov et al. 1986; Squyres et al., 1992, see Fig. 8). This term is not an established geologic one. Although it may be useful in certain circumstances, using more general terms such as “mountain ranges” may better align the study of Venus geology with the other terrestrial worlds (Klimczak et al., 2019). In any
<figure><img src="image_10.png" /><figcaption>**Fig. 8** - Left-look Magellan SAR showing radar-bright wrinkle ridges trending ~E-W with ~20 km spacing</figcaption></figure>

case, these belts are by any measure orogenic, in that they represent the (local) convergence of crustal material. Generally appearing as broad, linear rises a few hundred metres in relief, tens of kilometres in width, and many hundreds of kilometres long (Squyres et al., 1992), these systems typically show multiple anastomosing secondary arches and ridges superposed on the larger rise that superficially resemble the “wrinkle ridge” anticlinal folds commonly observed elsewhere on Venus and on the Moon, Mars, and Mercury. Larger mountain ranges have long been recognized on Venus, and given their sizes are generally referred to as such (e.g., Barsukov et al. 1986). Those wrinkle ridges are typically tens to hundreds of kilometres long but only a few pixels across (corresponding to a width of a few hundred metres) when seen in Magellan data. Venus’s wrinkle ridges are likely thrust-fault-related folds, given their morphological similarity to such features on other planetary bodies, which individually denote low amounts of shortening strain. These structures are ubiquitous on the planet, deforming the two most dominant surface units (the so-called upper and lower “ridged plains”) mapped by Bilotti and Suppe (1999) and Ivanov and Head (2011).

## 3.4 Polygons and Distributed Deformation

In addition to wrinkle ridges, numerous other tectonic features occur in the plains, comprising polygonal fractures and small-scale fracture patterns (Fig. 9). Large fields of polygonal fractures with spacing ranging from a ~1–2 km (essentially the limit of resolution) up to 25 km have been identified in over 200 regions (Moreels and Smrekar, 2003). These features have been proposed to form as a result of cooling of lava flows, analogous to columnar fractures in basaltic lava flows, although extremely thick flows would be required for these Venus features (Johnson and Sandwell, 1992). In contrast, the largest columnar joint patterns on Earth are up to 30 m. For this reason, as well as their frequent association with volcanic features, Johnson and Sandwell (1994) favoured a model with thermal stresses from either cooling lava flows or heating at the base of the lithosphere. An alternative explanation is the propagation of climate-change driven cooling into the subsurface (e.g., Anderson
<figure><img src="image_10.png" /><figcaption>**Fig. 8** - Left-look Magellan SAR showing radar-bright wrinkle ridges trending ~E-W with ~20 km spacing</figcaption></figure>

case, these belts are by any measure orogenic, in that they represent the (local) convergence of crustal material. Generally appearing as broad, linear rises a few hundred metres in relief, tens of kilometres in width, and many hundreds of kilometres long (Squyres et al., 1992), these systems typically show multiple anastomosing secondary arches and ridges superposed on the larger rise that superficially resemble the “wrinkle ridge” anticlinal folds commonly observed elsewhere on Venus and on the Moon, Mars, and Mercury. Larger mountain ranges have long been recognized on Venus, and given their sizes are generally referred to as such (e.g., Barsukov et al. 1986). Those wrinkle ridges are typically tens to hundreds of kilometres long but only a few pixels across (corresponding to a width of a few hundred metres) when seen in Magellan data. Venus’s wrinkle ridges are likely thrust-fault-related folds, given their morphological similarity to such features on other planetary bodies, which individually denote low amounts of shortening strain. These structures are ubiquitous on the planet, deforming the two most dominant surface units (the so-called upper and lower “ridged plains”) mapped by Bilotti and Suppe (1999) and Ivanov and Head (2011).

## 3.4 Polygons and Distributed Deformation

In addition to wrinkle ridges, numerous other tectonic features occur in the plains, comprising polygonal fractures and small-scale fracture patterns (Fig. 9). Large fields of polygonal fractures with spacing ranging from a ~1–2 km (essentially the limit of resolution) up to 25 km have been identified in over 200 regions (Moreels and Smrekar, 2003). These features have been proposed to form as a result of cooling of lava flows, analogous to columnar fractures in basaltic lava flows, although extremely thick flows would be required for these Venus features (Johnson and Sandwell, 1992). In contrast, the largest columnar joint patterns on Earth are up to 30 m. For this reason, as well as their frequent association with volcanic features, Johnson and Sandwell (1994) favoured a model with thermal stresses from either cooling lava flows or heating at the base of the lithosphere. An alternative explanation is the propagation of climate-change driven cooling into the subsurface (e.g., Anderson
<figure><img src="image_11.png" /><figcaption>**Fig. 9** - This section of FMAP 37N007 (Fig. 4 from Smrekar et al. 2002), shows two scales of polygons. The larger ones are controlled by intersections of NW-SE-trending wrinkle ridges</figcaption></figure>

and Smrekar, 1999; Solomon et al., 1999; Smrekar et al. 2002). This hypothesis relies on huge volumes of volcanism, as predicted by the catastrophic resurfacing hypothesis, and associated outgassing as a driver for hundreds of degrees of climate change (Bullock and Grinspoon 1996; 2001). These two particular models differ in detail: Anderson and Smrekar (1999) used a continuous-plate model of deformation due to thermal stresses to assess the effect of large predicted positive and negative temperature changes, and found that thermal cooling stresses can predict the observed polygonal features, but that heating is insufficient to produce deformation. In contrast, Solomon et al. (1999) employed a broken-plate model, as had been used for mid-ocean ridge studies, and found that, in that configuration, heating can produce wrinkle ridges and cooling can produce polygonal fractures. The origin of these features may be tested by determining the minimum size of the polygons with higher-resolution imaging, and whether or not they occur as flows or other distinct topographic features using high resolution topography.

## 3.5 Subduction and Terrestrial-Style Plate Tectonics

Subduction is believed to be a necessary first step for plate tectonics (e.g., Whattam and Stern, 2015; Lithgow-Bertelloni and Richards, 1995). Since plate tectonics began long ago
<figure><img src="image_11.png" /><figcaption>**Fig. 9** - This section of FMAP 37N007 (Fig. 4 from Smrekar et al. 2002), shows two scales of polygons. The larger ones are controlled by intersections of NW-SE-trending wrinkle ridges</figcaption></figure>

and Smrekar, 1999; Solomon et al., 1999; Smrekar et al. 2002). This hypothesis relies on huge volumes of volcanism, as predicted by the catastrophic resurfacing hypothesis, and associated outgassing as a driver for hundreds of degrees of climate change (Bullock and Grinspoon 1996; 2001). These two particular models differ in detail: Anderson and Smrekar (1999) used a continuous-plate model of deformation due to thermal stresses to assess the effect of large predicted positive and negative temperature changes, and found that thermal cooling stresses can predict the observed polygonal features, but that heating is insufficient to produce deformation. In contrast, Solomon et al. (1999) employed a broken-plate model, as had been used for mid-ocean ridge studies, and found that, in that configuration, heating can produce wrinkle ridges and cooling can produce polygonal fractures. The origin of these features may be tested by determining the minimum size of the polygons with higher-resolution imaging, and whether or not they occur as flows or other distinct topographic features using high resolution topography.

## 3.5 Subduction and Terrestrial-Style Plate Tectonics

Subduction is believed to be a necessary first step for plate tectonics (e.g., Whattam and Stern, 2015; Lithgow-Bertelloni and Richards, 1995). Since plate tectonics began long ago
on Earth (~4–1 Ga, still debated), neither the mechanism for initiation of subduction nor for plate tectonics is clearly identifiable. Thus, understanding the nature of subduction on Venus and the conditions that enable it to form would provide important insights into the initiation of terrestrial subduction. In addition to creating the slab-pull force that helps drive terrestrial plate tectonics, subduction plays a key role in recycling volatiles into the interior. These recycled volatiles can later be released back into the atmosphere, stabilizing Earth’s climate. On Venus, subduction provides links to interior processes and contributes to heat loss (e.g., Gülcher et al. 2020, 2023), and a means to estimate local elastic and lithospheric thickness.

Numerous locations (e.g., Artemis Corona, Fig. 10) on Venus appear to exhibit ‘rollback’ subduction, a process in which a plate sinks into the mantle under some combination of surface loading and negative plate buoyancy, without requiring lateral plate motion. The plate ‘rolls back’ as the downward-flexed plate advances away from the original subduction location as the plate sinks further. By analogy with terrestrial subduction zones such as the arcuate trenches of the South Sandwich Islands, the Aleutian trench, and north Fiji Basin, McKenzie et al. (1992), Sandwell and Schubert (1992a, 1992b), and Schubert and Sandwell (1995) identified a dozen sites of possible subduction totalling 10,000 km in length. Their criteria included a deep, narrow trench with arcuate or linear planform (Fig. 10), a curvature >10⁻⁷ m⁻¹ on the ‘outer rise’ side of the trench (the side attached to the subducting plate, e.g., on the southeast, outer side of Artemis), fractures parallel to the strike of the trench on both sides, and the absence of fractures cutting across the trenches. Hansen and Phillips (1993) presented an alternative interpretation of some of these features as resulting from mantle upwellings. This interpretation was based on the arcuate shape of many of these features, abundant volcanism in many locations, and fractures crossing the trench in one location.

Even with those examples of possible roll-back subduction associated with coronae, there is no evidence for a global network of tectonic plates diving into the planet’s interior or being produced anew at spreading centres (e.g., Solomon et al. 1992). Harris and Bédard (2014, 2015) provided a general overview of strike-slip faulting on Venus. They argued for substantial lateral motions and indenter-like escape tectonics in Ishtar Terra, hypothesising that Lakshmi Planum collided with Ishtar Terra. These authors further suggested the presence of lateral displacements of hundreds to thousands of kilometres within shear zones in Ovda and Thetis Regiones. Harris and Bédard (2014, 2015) proposed that mantle flow tractions could have acted on the deep lithospheric keels of Lakshmi Planum and Ovda and Thetis Regiones, driving them across the planet in a manner akin to how large continental blocks move on Earth. Importantly, such motions produce the kinds of shortening and transpressional structures, and at the same approximate scales, as are recorded in the Archean rocks. If this hypothesis is correct, then the large-scale “drift” of major terranes on Venus may be a contemporary analogue of Archean Earth, when cratonic nuclei were mobile but modern plate tectonics, including subduction, had yet to take hold (Harris and Bédard 2014, 2015).

# 4 Crust, Lithosphere and Heat Flow

The surface geologic processes described above are a result of the interior heat loss mechanisms of convection and conduction. The crust, a compositional layer, and the lithosphere, a mechanical layer, strongly control the specific manifestation of these geologic processes. Models of both geologic processes and interior evolution require assumptions about crustal and lithospheric properties. In addition, viscous-deformation models require knowledge of
on Earth (~4–1 Ga, still debated), neither the mechanism for initiation of subduction nor for plate tectonics is clearly identifiable. Thus, understanding the nature of subduction on Venus and the conditions that enable it to form would provide important insights into the initiation of terrestrial subduction. In addition to creating the slab-pull force that helps drive terrestrial plate tectonics, subduction plays a key role in recycling volatiles into the interior. These recycled volatiles can later be released back into the atmosphere, stabilizing Earth’s climate. On Venus, subduction provides links to interior processes and contributes to heat loss (e.g., Gülcher et al. 2020, 2023), and a means to estimate local elastic and lithospheric thickness.

Numerous locations (e.g., Artemis Corona, Fig. 10) on Venus appear to exhibit ‘rollback’ subduction, a process in which a plate sinks into the mantle under some combination of surface loading and negative plate buoyancy, without requiring lateral plate motion. The plate ‘rolls back’ as the downward-flexed plate advances away from the original subduction location as the plate sinks further. By analogy with terrestrial subduction zones such as the arcuate trenches of the South Sandwich Islands, the Aleutian trench, and north Fiji Basin, McKenzie et al. (1992), Sandwell and Schubert (1992a, 1992b), and Schubert and Sandwell (1995) identified a dozen sites of possible subduction totalling 10,000 km in length. Their criteria included a deep, narrow trench with arcuate or linear planform (Fig. 10), a curvature >10⁻⁷ m⁻¹ on the ‘outer rise’ side of the trench (the side attached to the subducting plate, e.g., on the southeast, outer side of Artemis), fractures parallel to the strike of the trench on both sides, and the absence of fractures cutting across the trenches. Hansen and Phillips (1993) presented an alternative interpretation of some of these features as resulting from mantle upwellings. This interpretation was based on the arcuate shape of many of these features, abundant volcanism in many locations, and fractures crossing the trench in one location.

Even with those examples of possible roll-back subduction associated with coronae, there is no evidence for a global network of tectonic plates diving into the planet’s interior or being produced anew at spreading centres (e.g., Solomon et al. 1992). Harris and Bédard (2014, 2015) provided a general overview of strike-slip faulting on Venus. They argued for substantial lateral motions and indenter-like escape tectonics in Ishtar Terra, hypothesising that Lakshmi Planum collided with Ishtar Terra. These authors further suggested the presence of lateral displacements of hundreds to thousands of kilometres within shear zones in Ovda and Thetis Regiones. Harris and Bédard (2014, 2015) proposed that mantle flow tractions could have acted on the deep lithospheric keels of Lakshmi Planum and Ovda and Thetis Regiones, driving them across the planet in a manner akin to how large continental blocks move on Earth. Importantly, such motions produce the kinds of shortening and transpressional structures, and at the same approximate scales, as are recorded in the Archean rocks. If this hypothesis is correct, then the large-scale “drift” of major terranes on Venus may be a contemporary analogue of Archean Earth, when cratonic nuclei were mobile but modern plate tectonics, including subduction, had yet to take hold (Harris and Bédard 2014, 2015).

# 4 Crust, Lithosphere and Heat Flow

The surface geologic processes described above are a result of the interior heat loss mechanisms of convection and conduction. The crust, a compositional layer, and the lithosphere, a mechanical layer, strongly control the specific manifestation of these geologic processes. Models of both geologic processes and interior evolution require assumptions about crustal and lithospheric properties. In addition, viscous-deformation models require knowledge of
<figure><img src="image_12.png" /><figcaption>**Fig. 10** - Magellan image data overlain on global topography of Artemis Corona. Artemis is the largest of all coronae, with a diameter of ~2600 km. Its southeast is a likely site of subduction. Relative topography is in m</figcaption></figure>

strain rate since the material strength is a function of the rate at which stress is applied. Models of deformation require laboratory-derived relationships between stress and strain. Even for Earth studies, there are substantial open questions about how to apply these laboratory-derived rheological laws to geologic processes, with laboratory predictions, done at very high strain rates of necessity, apparently overpredicting the strength of rocks deforming at the lower but more geologically realistic strain rates. In this section, we describe the available data to constrain estimates of Venus' crustal and lithospheric rheology, and review what is inferred about the composition and thickness of the crust, the temperature and thickness of the lithosphere, and strain rates based on terrestrial analogy.

## 4.1 Crust and Upper Mantle

### 4.1.1 Crust and Surface Composition from Geochemical Constraints

The similar bulk densities of Venus and Earth imply that they have broadly similar elemental contents (e.g., Taylor and McLennan, 1995; Lee et al. 2009; Rubie et al., 2015; Weller and
<figure><img src="image_12.png" /><figcaption>**Fig. 10** - Magellan image data overlain on global topography of Artemis Corona. Artemis is the largest of all coronae, with a diameter of ~2600 km. Its southeast is a likely site of subduction. Relative topography is in m</figcaption></figure>

strain rate since the material strength is a function of the rate at which stress is applied. Models of deformation require laboratory-derived relationships between stress and strain. Even for Earth studies, there are substantial open questions about how to apply these laboratory-derived rheological laws to geologic processes, with laboratory predictions, done at very high strain rates of necessity, apparently overpredicting the strength of rocks deforming at the lower but more geologically realistic strain rates. In this section, we describe the available data to constrain estimates of Venus' crustal and lithospheric rheology, and review what is inferred about the composition and thickness of the crust, the temperature and thickness of the lithosphere, and strain rates based on terrestrial analogy.

## 4.1 Crust and Upper Mantle

### 4.1.1 Crust and Surface Composition from Geochemical Constraints

The similar bulk densities of Venus and Earth imply that they have broadly similar elemental contents (e.g., Taylor and McLennan, 1995; Lee et al. 2009; Rubie et al., 2015; Weller and
Duncan 2015; Shellnutt 2016). However, the crusts of Venus and Earth could have quite different, and variable, compositions. In-situ data on crustal composition is limited to data from several Soviet landers in plains locations with instrumentation that measured elemental geochemistry (see Treiman 2007 for a review). Measured compositions in these locations were primarily mafic and likely basaltic (Surkov et al. 1984, 1986; Kargel et al. 1993). The rocks for which near-complete major elemental chemistry measurements were made are terrestrial equivalents of tholeiitic (Venera 14) and alkalic (Venera 13) basalt, whereas the Venera 8 rock could be similar (e.g., Kargel et al. 1993) or notably different, with anomalously high K, Th, and U concentrations similar to terrestrial granodiorite or dacite (e.g., Nikolayeva 1990). Trace elemental (K, Th, U) analyses at other landing sites (Vega 1, Venera 9, Venera 10) also indicated the presence of terrestrial basalt. On the basis of morphology, much of Venus is inferred to be covered in basaltic material, similar to the measured plains locations.

There is some evidence from flyby and orbital spectroscopy for felsic rock in the massive tessera plateaus (e.g., Hashimoto et al. 2008; Helbert et al., 2008; Mueller et al. 2008; Gilmore et al. 2015; Gilmore et al., 2023). A small suite of features such as 'festoons' also have morphologies arguably consistent with high viscosity silicious lavas (e.g., Pavri et al. 1992; Bridges 1997), but perhaps also with basaltic lavas (e.g., Stofan et al., 2000; Wroblewski et al. 2019). However, morphology is a poor indicator of composition and in no way a substitute for remote or in-situ chemical data. Thus, the degree and scale of crustal differentiation that has taken place on Venus is debated, with some arguing for less than Earth's continents even in the tesserae and festoon flows (e.g., Kargel et al. 1993; Grimm and Hess, 1997; Treiman 2007; Treiman et al. 2016; Wroblewski et al., 2019).

Magma differentiation processes in terrestrial intraplate settings, thought to be the closest analogue for Venus tectonic settings (e.g., rifts, hotspots), are primarily related to crystallisation (i.e., equilibrium or fractional) and partial melting. In contrast, mixing or mingling is of secondary influence in comparison to convergent margin settings (Bachman and Bergantz, 2008; Christiansen and McCurry, 2008). Thus, forward petrological modelling using a range of magmatic conditions (i.e., water content, oxygen fugacity, and pressure), together with the rock compositions at the Venera 13 and Venera 14 landing sites and to a lesser extent the Vega 2 compositional measurements, can offer useful insight into the possible igneous rocks that could comprise a portion of the upper to middle crust. Fractional crystallisation modelling demonstrates that melt compositions similar to within-plate intermediate and silicic rocks (e.g., dacite, granodiorite, andesite, diorite, rhyolite, granite, trachyte, syenite), including the theoretical whole-rock composition of the Venera 8 landing site rocks and their K, Th, and U concentrations, can be produced at middle-to-upper crustal pressures (i.e., ~0.1–0.5 GPa) under reducing or oxidizing, and hydrous or anhydrous, conditions (Shellnutt 2013, 2018, 2019). The residual solid after fractionation is mostly composed of cumulus olivine, orthopyroxene, clinopyroxene, plagioclase, and Fe–Ti oxide minerals (i.e., titanomagnetite, ilmenite) and are analogous to the cumulate rocks of layered mafic–ultramafic intrusions of Earth, such as the Bushveld Complex, Stillwater Intrusion, and Skaergaard Intrusion. On Earth, rocks similar to the modelled residual silicic liquids and their mafic/ultramafic cumulate rocks are commonly found at intraplate settings associated with rifting and hotspot volcanism. Therefore, they are likely present within the crust of Venus as layered igneous complexes, and their silicic volcanic rocks may be associated with volcanic rises, coronae, pancake domes, and shield volcanoes. The lower viscosity of intraplate silicic volcanic rocks compared with convergent margin silicic volcanic rocks suggests that their volcanic edifices may not necessarily have meaningfully different morphologies from their mafic counterparts (Christiansen, 2005). As discussed above, volcanic morphology is suggestive of composition (e.g., Wroblewski et al. 2019), but not necessarily definitively so.
Duncan 2015; Shellnutt 2016). However, the crusts of Venus and Earth could have quite different, and variable, compositions. In-situ data on crustal composition is limited to data from several Soviet landers in plains locations with instrumentation that measured elemental geochemistry (see Treiman 2007 for a review). Measured compositions in these locations were primarily mafic and likely basaltic (Surkov et al. 1984, 1986; Kargel et al. 1993). The rocks for which near-complete major elemental chemistry measurements were made are terrestrial equivalents of tholeiitic (Venera 14) and alkalic (Venera 13) basalt, whereas the Venera 8 rock could be similar (e.g., Kargel et al. 1993) or notably different, with anomalously high K, Th, and U concentrations similar to terrestrial granodiorite or dacite (e.g., Nikolayeva 1990). Trace elemental (K, Th, U) analyses at other landing sites (Vega 1, Venera 9, Venera 10) also indicated the presence of terrestrial basalt. On the basis of morphology, much of Venus is inferred to be covered in basaltic material, similar to the measured plains locations.

There is some evidence from flyby and orbital spectroscopy for felsic rock in the massive tessera plateaus (e.g., Hashimoto et al. 2008; Helbert et al., 2008; Mueller et al. 2008; Gilmore et al. 2015; Gilmore et al., 2023). A small suite of features such as 'festoons' also have morphologies arguably consistent with high viscosity silicious lavas (e.g., Pavri et al. 1992; Bridges 1997), but perhaps also with basaltic lavas (e.g., Stofan et al., 2000; Wrobleski et al. 2019). However, morphology is a poor indicator of composition and in no way a substitute for remote or in-situ chemical data. Thus, the degree and scale of crustal differentiation that has taken place on Venus is debated, with some arguing for less than Earth's continents even in the tesserae and festoon flows (e.g., Kargel et al. 1993; Grimm and Hess, 1997; Treiman 2007; Treiman et al. 2016; Wrobleski et al., 2019).

Magma differentiation processes in terrestrial intraplate settings, thought to be the closest analogue for Venus tectonic settings (e.g., rifts, hotspots), are primarily related to crystallisation (i.e., equilibrium or fractional) and partial melting. In contrast, mixing or mingling is of secondary influence in comparison to convergent margin settings (Bachman and Bergantz, 2008; Christiansen and McCurry, 2008). Thus, forward petrological modelling using a range of magmatic conditions (i.e., water content, oxygen fugacity, and pressure), together with the rock compositions at the Venera 13 and Venera 14 landing sites and to a lesser extent the Vega 2 compositional measurements, can offer useful insight into the possible igneous rocks that could comprise a portion of the upper to middle crust. Fractional crystallisation modelling demonstrates that melt compositions similar to within-plate intermediate and silicic rocks (e.g., dacite, granodiorite, andesite, diorite, rhyolite, granite, trachyte, syenite), including the theoretical whole-rock composition of the Venera 8 landing site rocks and their K, Th, and U concentrations, can be produced at middle-to-upper crustal pressures (i.e., ~0.1–0.5 GPa) under reducing or oxidizing, and hydrous or anhydrous, conditions (Shellnutt 2013, 2018, 2019). The residual solid after fractionation is mostly composed of cumulus olivine, orthopyroxene, clinopyroxene, plagioclase, and Fe–Ti oxide minerals (i.e., titanomagnetite, ilmenite) and are analogous to the cumulate rocks of layered mafic–ultramafic intrusions of Earth, such as the Bushveld Complex, Stillwater Intrusion, and Skaergaard Intrusion. On Earth, rocks similar to the modelled residual silicic liquids and their mafic/ultramafic cumulate rocks are commonly found at intraplate settings associated with rifting and hotspot volcanism. Therefore, they are likely present within the crust of Venus as layered igneous complexes, and their silicic volcanic rocks may be associated with volcanic rises, coronae, pancake domes, and shield volcanoes. The lower viscosity of intraplate silicic volcanic rocks compared with convergent margin silicic volcanic rocks suggests that their volcanic edifices may not necessarily have meaningfully different morphologies from their mafic counterparts (Christiansen, 2005). As discussed above, volcanic morphology is suggestive of composition (e.g., Wrobleski et al. 2019), but not necessarily definitively so.
Partial melting of the mafic crust of Venus can also yield magmas with intermediate-to-silicic compositions under high pressure and temperature (~1 GPa, >1250 K) conditions expected for the lower crust of Venus (James et al. 2013; Shellnutt 2013). Of particular interest is the Venera 14 landing site, which could be very similar to basalts of terrestrial Archean greenstone belts. It is interesting because rocks of similar composition to those near Venera 14 can yield silicic liquids that are indistinguishable from the tonalite-trondhjemite-granodiorite (TTG) series of rocks found in Archean cratons (Shellnutt 2013, 2018; Johnson et al., 2017). The formation of some TTGs may be related to thickening of mafic crust and the subsequent partial melting of the lower parts of the thickened crust either by ambient geothermal conditions or by anomalously hot mantle plumes (Smithies, 2000; Moyen and Martin, 2012). Venus has regions of tectonically thickened crust and evidence for hotspot volcanism and high temperature, and low viscosity (perhaps ultramafic) lava flows (e.g., Lenardic et al., 1991; Komatsu and Baker 1994; Hansen et al., 1999) that may be experiencing similar processes. If TTG suites are generated on Venus, they might form within regions of thickened crust such as Ishtar Terra and Ovda Regio. Additionally, gravitational instabilities, in the form of rollback subduction, delamination, or lithospheric dripping (see Sect. 3.5. above), may occur on Venus. Such instabilities would lead to remelting of the crust, recycling of volatiles, and potentially to the production of more evolved lavas erupting at the surface, including even low-viscosity flows (Elkins-Tanton et al. 2007).

The basalt at the Venera 14 site is broadly similar to olivine tholeiite and possibly normal-mid-ocean ridge basalt (N-MORB) and thus may have originated from a mantle thermal regime similar to Earth (McKenzie et al., 1992; Kargel et al. 1993; Fegly Jr., 2003; Filiberto, 2014). The material measured at the Vega 2 landing site appears to be a mixture of soil or weathered material and rock, since the composition is broadly basaltic but the SO₃ content is extremely high (~4.7 ± 1.5 wt%). If those measurements were accurate, such rocks would not have a direct terrestrial equivalent, with the exception of a mixture of sulphate rock (e.g., gypsum, anhydrite) or sulphide ore with a mafic (basalt, gabbro) host rock that is typical of some basalt country rock relationships, layered intrusions, and/or volcanogenic mass-sulphide deposits (VMS). Sulphate rocks near Vega 2 might be related to the carbonate sulphate volcanism that was proposed to form the canali (e.g., Kargel et al. 1994).

## 4.1.2 Crustal Thickness and Upper Mantle Viscosity

Gravity and topography measurements are currently the only datasets available with which we can investigate the interior structure of Venus. Without more direct measurements, derivation of interior structure is inherently non-unique and subject to the assumptions made about such parameters as crust and mantle density, thermal gradient, and strain rate. A key challenge for Venus is the resolution of the measured gravity field, which ranges from ~400 to ~1000 km (Konopliv et al., 1999). The spherical harmonic field is expanded to degree and order 180 to prevent aliasing of the signal. The actual resolution of the field is represented by degree strength, which is a maximum of degree and order ~100 (Konopliv et al., 1999).

One approach to interpreting gravity and topography is to invert their spherical harmonics representation to provide an estimate of crustal thickness and upper-mantle viscosity and temperature (e.g., Banerdt, 1986; Herrick and Phillips, 1992; James et al. 2013). The correlations between thick/thin crust and hot/cold upper mantle with geologic surface features provides insight into the relationship between mantle convection and volcanism and tectonics. Different lithospheric properties and upper-mantle viscosity structures can be assumed in the inversions and the results then evaluated in terms of realism (e.g., are the temperature
Partial melting of the mafic crust of Venus can also yield magmas with intermediate-to-silicic compositions under high pressure and temperature (~1 GPa, >1250 K) conditions expected for the lower crust of Venus (James et al. 2013; Shellnutt 2013). Of particular interest is the Venera 14 landing site, which could be very similar to basalts of terrestrial Archean greenstone belts. It is interesting because rocks of similar composition to those near Venera 14 can yield silicic liquids that are indistinguishable from the tonalite-trondhjemite-granodiorite (TTG) series of rocks found in Archean cratons (Shellnutt 2013, 2018; Johnson et al., 2017). The formation of some TTGs may be related to thickening of mafic crust and the subsequent partial melting of the lower parts of the thickened crust either by ambient geothermal conditions or by anomalously hot mantle plumes (Smithies, 2000; Moyen and Martin, 2012). Venus has regions of tectonically thickened crust and evidence for hotspot volcanism and high temperature, and low viscosity (perhaps ultramafic) lava flows (e.g., Lenardic et al., 1991; Komatsu and Baker 1994; Hansen et al., 1999) that may be experiencing similar processes. If TTG suites are generated on Venus, they might form within regions of thickened crust such as Ishtar Terra and Ovda Regio. Additionally, gravitational instabilities, in the form of rollback subduction, delamination, or lithospheric dripping (see Sect. 3.5. above), may occur on Venus. Such instabilities would lead to remelting of the crust, recycling of volatiles, and potentially to the production of more evolved lavas erupting at the surface, including even low-viscosity flows (Elkins-Tanton et al. 2007).

The basalt at the Venera 14 site is broadly similar to olivine tholeiite and possibly normal-mid-ocean ridge basalt (N-MORB) and thus may have originated from a mantle thermal regime similar to Earth (McKenzie et al., 1992; Kargel et al. 1993; Fegly Jr., 2003; Filiberto, 2014). The material measured at the Vega 2 landing site appears to be a mixture of soil or weathered material and rock, since the composition is broadly basaltic but the SO₃ content is extremely high (~4.7 ± 1.5 wt%). If those measurements were accurate, such rocks would not have a direct terrestrial equivalent, with the exception of a mixture of sulphate rock (e.g., gypsum, anhydrite) or sulphide ore with a mafic (basalt, gabbro) host rock that is typical of some basalt country rock relationships, layered intrusions, and/or volcanogenic mass-sulphide deposits (VMS). Sulphate rocks near Vega 2 might be related to the carbonate sulphate volcanism that was proposed to form the canali (e.g., Kargel et al. 1994).

## 4.1.2 Crustal Thickness and Upper Mantle Viscosity

Gravity and topography measurements are currently the only datasets available with which we can investigate the interior structure of Venus. Without more direct measurements, derivation of interior structure is inherently non-unique and subject to the assumptions made about such parameters as crust and mantle density, thermal gradient, and strain rate. A key challenge for Venus is the resolution of the measured gravity field, which ranges from ~400 to ~1000 km (Konopliv et al., 1999). The spherical harmonic field is expanded to degree and order 180 to prevent aliasing of the signal. The actual resolution of the field is represented by degree strength, which is a maximum of degree and order ~100 (Konopliv et al., 1999).

One approach to interpreting gravity and topography is to invert their spherical harmonics representation to provide an estimate of crustal thickness and upper-mantle viscosity and temperature (e.g., Banerdt, 1986; Herrick and Phillips, 1992; James et al. 2013). The correlations between thick/thin crust and hot/cold upper mantle with geologic surface features provides insight into the relationship between mantle convection and volcanism and tectonics. Different lithospheric properties and upper-mantle viscosity structures can be assumed in the inversions and the results then evaluated in terms of realism (e.g., are the temperature
anomalies consistent with expected mantle behaviour?). The relationship between gravity highs and lows and stress patterns has also been used to infer areas of present or recent activity (e.g. Smrekar, 1994). The long-wavelength geoid (the integral of the gravity field) can be interpreted as an anomalous density distribution that gives rise to lithospheric stresses (Sandwell et al., 1997).

Two-layer modelling with a low-order, pre-Magellan gravity field indicated that the mantle temperatures required to reproduce the observed gravity and topography become unrealistic if the viscosity structure includes an Earth-like asthenosphere (Herrick and Phillips, 1992). Other studies also suggested that Venus lacks a narrow, low-viscosity zone (e.g., Kiefer and Hager, 1991; Smrekar and Phillips 1991; Steinberger et al. 2010). However, different parameterizations (Pauer et al. 2006) and an approach of first removing large areas of thick crust (Maia et al. 2023) produced results that favor a ~200 km-thick zone below the lithosphere with a viscosity ~10–100 times lower than the mantle beneath, i.e., more akin to Earth’s asthenosphere.

A low-viscosity asthenosphere on Venus may indicate that the interior of Venus is wet. Importantly, however, alternative explanations are possible. A low-viscosity zone can be produced by CO₂ rather than water (Sifré et al., 2014; Ghail 2015; Chantel et al., 2016) and by a specific dry olivine activation volume (Armann and Tackley 2012).

The associations between crustal thickness and inferred mantle convection with geologic features (e.g., James et al. 2013) clearly show that areas of thickened crust are associated with tessera plateaus, while significant mantle upwellings are associated with large volcanic rises (Fig. 11), some of which are connected by major rift systems.

The overall pattern of surface features relative to calculated mantle convection and crustal thickness can be used to make some inferences regarding the evolution of the interior over time. The correlation of the present gravity field with the global volcano–rift–corona pattern in many regions suggest that the large-scale mantle convection pattern has not changed dramatically during the time period over which these features were emplaced. The global wrinkle ridge pattern generally, but not everywhere, roughly parallels the geoid gradient, consistent with these features forming in response to current mantle patterns (Bilotti and Suppe, 1999; Sandwell et al., 1997). Laufey Regio, which hosts large volcanoes and concentric wrinkle ridges, appears to have formed in response to a mantle upwelling that has waned or become inactive (Brian et al. 2004). That there is no clear association of tessera plateaus with the inferred mantle convection pattern is consistent with most, but not necessarily all – Ishtar Terra is a notable exception, of these areas being inactive. While they may no longer be forming, high topography is likely to be gravitationally relaxing, depending on its crustal strength (e.g., Smrekar and Phillips, 1988; Nunes et al. 2004; Romeo and Turcotte 2008; Maia and Wieczorek, 2022; Nimmo and Mackwell 2023).

Most estimates of Venus’ average crustal thickness are tens of km, with regional maxima up to ~70–90 km. Anderson and Smrekar (2006) and Jimenez-Daiz et al. (2015) both compiled global maps of crustal thickness using spectral admittance methods to solve for both local crustal and elastic thickness, without a contribution from mantle dynamics. Using somewhat different approaches, Anderson and Smrekar (2006) find an average crustal thickness on the order of ~10–20 km and Jimenez-Daiz et al. (2015) find ~20–25 km. James et al. (2013) employed a different method: analysing geoid to topography ratios and performing a two-layer layer inversion of crustal thicknesses and mantle forces to derive smoothed maps of global crustal thickness and dynamic mantle forces. They assumed a uniform elastic thickness of 20 km, which is reasonable (see the next subsection). They estimated a mean crustal thickness of ~8–25 km with the assumption that crustal thickness is non-zero everywhere and is limited by the transition of basalt to eclogite at ~60–70 km.
anomalies consistent with expected mantle behaviour?). The relationship between gravity highs and lows and stress patterns has also been used to infer areas of present or recent activity (e.g. Smrekar, 1994). The long-wavelength geoid (the integral of the gravity field) can be interpreted as an anomalous density distribution that gives rise to lithospheric stresses (Sandwell et al., 1997).

Two-layer modelling with a low-order, pre-Magellan gravity field indicated that the mantle temperatures required to reproduce the observed gravity and topography become unrealistic if the viscosity structure includes an Earth-like asthenosphere (Herrick and Phillips, 1992). Other studies also suggested that Venus lacks a narrow, low-viscosity zone (e.g., Kiefer and Hager, 1991; Smrekar and Phillips 1991; Steinberger et al. 2010). However, different parameterizations (Pauer et al. 2006) and an approach of first removing large areas of thick crust (Maia et al. 2023) produced results that favor a ~200 km-thick zone below the lithosphere with a viscosity ~10–100 times lower than the mantle beneath, i.e., more akin to Earth's asthenosphere.

A low-viscosity asthenosphere on Venus may indicate that the interior of Venus is wet. Importantly, however, alternative explanations are possible. A low-viscosity zone can be produced by CO₂ rather than water (Sifré et al., 2014; Ghail 2015; Chantel et al., 2016) and by a specific dry olivine activation volume (Armann and Tackley 2012).

The associations between crustal thickness and inferred mantle convection with geologic features (e.g., James et al. 2013) clearly show that areas of thickened crust are associated with tessera plateaus, while significant mantle upwellings are associated with large volcanic rises (Fig. 11), some of which are connected by major rift systems.

The overall pattern of surface features relative to calculated mantle convection and crustal thickness can be used to make some inferences regarding the evolution of the interior over time. The correlation of the present gravity field with the global volcano–rift–corona pattern in many regions suggest that the large-scale mantle convection pattern has not changed dramatically during the time period over which these features were emplaced. The global wrinkle ridge pattern generally, but not everywhere, roughly parallels the geoid gradient, consistent with these features forming in response to current mantle patterns (Bilotti and Suppe, 1999; Sandwell et al., 1997). Laufey Regio, which hosts large volcanoes and concentric wrinkle ridges, appears to have formed in response to a mantle upwelling that has waned or become inactive (Brian et al. 2004). That there is no clear association of tessera plateaus with the inferred mantle convection pattern is consistent with most, but not necessarily all – Ishtar Terra is a notable exception, of these areas being inactive. While they may no longer be forming, high topography is likely to be gravitationally relaxing, depending on its crustal strength (e.g., Smrekar and Phillips, 1988; Nunes et al. 2004; Romeo and Turcotte 2008; Maia and Wieczorek, 2022; Nimmo and Mackwell 2023).

Most estimates of Venus' average crustal thickness are tens of km, with regional maxima up to ~70–90 km. Anderson and Smrekar (2006) and Jimenez-Daiz et al. (2015) both compiled global maps of crustal thickness using spectral admittance methods to solve for both local crustal and elastic thickness, without a contribution from mantle dynamics. Using somewhat different approaches, Anderson and Smrekar (2006) find an average crustal thickness on the order of ~10–20 km and Jimenez-Daiz et al. (2015) find ~20–25 km. James et al. (2013) employed a different method: analysing geoid to topography ratios and performing a two-layer layer inversion of crustal thicknesses and mantle forces to derive smoothed maps of global crustal thickness and dynamic mantle forces. They assumed a uniform elastic thickness of 20 km, which is reasonable (see the next subsection). They estimated a mean crustal thickness of ~8–25 km with the assumption that crustal thickness is non-zero everywhere and is limited by the transition of basalt to eclogite at ~60–70 km.
<figure><img src="image_13.png" /><figcaption>**Fig. 11** Two-layer solution of James et al. (2013) superposed on Cycle 1 Magellan imaging of Venus. Top panel shows crustal thickness in km, and bottom panel density contrast in the uppermost mantle in kg m⁻². In the bottom panel, these Regios have been identified as hotspots: Imdr, Atla, Beta, Eistla (Western, Central, Eastern), Themis, Bell, Dione and Lada Terra (Stofan et al. 1995). Red dots show locations of large volcanoes in the database of Crumpler et al. (1997) with diameters >500 km. Among the many relationships between gravity, topography, and geology observable in these maps, large volcanoes are most often found in association with mantle upwellings or major rift zones and mostly avoid the areas of thickest crust, particularly tessera regions</figcaption></figure>

Their calculated mean would increase to ~45 km if the Maxwell Montes formed tectonically in relatively recent times—meaning that they can be excluded from the assumption that crustal thickness cannot exceed ~70 km because of the basalt-eclogite phase transition has not reached completion, and/or the eclogite has not delaminated (e.g., Namiki and Solomon 1993; Jull and Arkani-Hamed, 1995).

Relatedly, convection models use gravity and topography to estimate a combination of crustal thickness and mantle properties (see Rolf et al. 2022, this collection, for more de-
<figure><img src="image_13.png" /><figcaption>**Fig. 11** Two-layer solution of James et al. (2013) superposed on Cycle 1 Magellan imaging of Venus. Top panel shows crustal thickness in km, and bottom panel density contrast in the uppermost mantle in kg m⁻². In the bottom panel, these Regios have been identified as hotspots: Imdr, Atla, Beta, Eistla (Western, Central, Eastern), Themis, Bell, Dione and Lada Terra (Stofan et al. 1995). Red dots show locations of large volcanoes in the database of Crumpler et al. (1997) with diameters >500 km. Among the many relationships between gravity, topography, and geology observable in these maps, large volcanoes are most often found in association with mantle upwellings or major rift zones and mostly avoid the areas of thickest crust, particularly tessera regions</figcaption></figure>

Their calculated mean would increase to ~45 km if the Maxwell Montes formed tectonically in relatively recent times—meaning that they can be excluded from the assumption that crustal thickness cannot exceed ~70 km because of the basalt-eclogite phase transition has not reached completion, and/or the eclogite has not delaminated (e.g., Namiki and Solomon 1993; Jull and Arkani-Hamed, 1995).

Relatedly, convection models use gravity and topography to estimate a combination of crustal thickness and mantle properties (see Rolf et al. 2022, this collection, for more de-
tails). Such models typically focus on the longest-wavelength portion of the gravity spec-
trum, and tend to yield larger estimates of crustal thickness. Steinberger et al. (2010) found
an average thickness of ~60 km using the range of the observed gravity spectrum that
is likely dominated by crustal sources (i.e., above spherical harmonic degree 40). Wei et
al. (2014) used models of Venus' mantle convection to include dynamical (mantle flow-
sourced) contributions to derive a crustal thickness range of ~28-70 km, with the greatest
values under highland regions such as Ishtar Terra and Aphrodite Terra.

## 4.2 Lithospheric Thickness and Heat Flow

A planet's lithosphere is the outermost part of the solid body. We can distinguish between three definitions of the lithosphere. First, the elastic lithosphere is the calculated thickness of a layer that is assumed to exhibit only elastic behavior. Second, the mechanical lithosphere represents a rheological layer that is too strong to convect even over long timescales. It encompasses the elastic lithosphere, and can additionally support loads via viscous deformation. Third, the thermal lithosphere contains the mechanical lithosphere and is defined as the layer in which heat is transported by thermal conduction, not convection. The thermal lithospheric thickness is usually defined by the depth to the mantle convection temperature, often taken as ~1575 K. In contrast, the crust and the uppermost mantle layers are defined by their composition. Composition implies a specific rheology, which has implications for the strength of the viscous portion of the mechanical lithosphere. The boundary between the crust and the uppermost mantle (or lithospheric mantle) may fall anywhere within these three layers, depending on parameters such as thermal gradient, strain rate, and rheology of the crust versus uppermost mantle.

The thermal lithosphere is essential to models of Venus's evolution. As the boundary layer, or 'lid,' of the convecting mantle, it couples the thermal state of the convecting interior to the surface. The heat transport across the thermal lithosphere limits the amount of heat leaving the mantle, determined by the thermal conductivity of mantle rocks, as well as the thickness and the temperature contrast across the lithosphere (i.e., Fourier's law). Both the temperature contrast and the thickness strongly depend on Venus' thermal history and its geodynamic regime (e.g., Rolf et al. 2022, this collection). The present thermal lithospheric thickness thus provides useful constraints on estimates of the thermal state of Venus' interior and how it may have evolved to this state.

Although convection models directly predict the thickness of the thermal lithosphere, it is the elastic lithospheric thickness that we can estimate more accurately. The wavelength at which the elastic lithosphere deforms in response to loads depends on its thickness. As discussed above, observations of gravity and topography thus yield quantitative estimates of the thickness of the elastic lithosphere. In idealized elastic models, lithospheric bending gives rise to internal stresses that increase linearly from zero at the mid-point of the elastic lithosphere and reach maximal values at its upper and lower boundaries. However, in reality, brittle failure and ductile flow limit the maximum possible stresses at the top and bottom of the mechanical lithosphere, respectively. Although the bottom of the mechanical lithosphere is viscous, it deforms only slowly and thus can also support surface topography or subsurface loads.

Specifically, the mechanical lithosphere encompasses both elastically and viscously strong parts, as defined by a yield-stress envelope. For a given rheological model of the lithosphere, the elastic thickness can be used to determine the mechanical thickness, which in turn is tied to the lithosphere's thermal state (e.g., McNutt 1984). The mechanical lithosphere is always thicker than the elastic lithosphere—but the total bending moments of both
tails). Such models typically focus on the longest-wavelength portion of the gravity spec-
trum, and tend to yield larger estimates of crustal thickness. Steinberger et al. (2010) found
an average thickness of ~60 km using the range of the observed gravity spectrum that
is likely dominated by crustal sources (i.e., above spherical harmonic degree 40). Wei et
al. (2014) used models of Venus' mantle convection to include dynamical (mantle flow-
sourced) contributions to derive a crustal thickness range of ~28-70 km, with the greatest
values under highland regions such as Ishtar Terra and Aphrodite Terra.

## 4.2 Lithospheric Thickness and Heat Flow

A planet's lithosphere is the outermost part of the solid body. We can distinguish between three definitions of the lithosphere. First, the elastic lithosphere is the calculated thickness of a layer that is assumed to exhibit only elastic behavior. Second, the mechanical lithosphere represents a rheological layer that is too strong to convect even over long timescales. It encompasses the elastic lithosphere, and can additionally support loads via viscous deformation. Third, the thermal lithosphere contains the mechanical lithosphere and is defined as the layer in which heat is transported by thermal conduction, not convection. The thermal lithospheric thickness is usually defined by the depth to the mantle convection temperature, often taken as ~1575 K. In contrast, the crust and the uppermost mantle layers are defined by their composition. Composition implies a specific rheology, which has implications for the strength of the viscous portion of the mechanical lithosphere. The boundary between the crust and the uppermost mantle (or lithospheric mantle) may fall anywhere within these three layers, depending on parameters such as thermal gradient, strain rate, and rheology of the crust versus uppermost mantle.

The thermal lithosphere is essential to models of Venus's evolution. As the boundary layer, or 'lid,' of the convecting mantle, it couples the thermal state of the convecting interior to the surface. The heat transport across the thermal lithosphere limits the amount of heat leaving the mantle, determined by the thermal conductivity of mantle rocks, as well as the thickness and the temperature contrast across the lithosphere (i.e., Fourier's law). Both the temperature contrast and the thickness strongly depend on Venus' thermal history and its geodynamic regime (e.g., Rolf et al. 2022, this collection). The present thermal lithospheric thickness thus provides useful constraints on estimates of the thermal state of Venus' interior and how it may have evolved to this state.

Although convection models directly predict the thickness of the thermal lithosphere, it is the elastic lithospheric thickness that we can estimate more accurately. The wavelength at which the elastic lithosphere deforms in response to loads depends on its thickness. As discussed above, observations of gravity and topography thus yield quantitative estimates of the thickness of the elastic lithosphere. In idealized elastic models, lithospheric bending gives rise to internal stresses that increase linearly from zero at the mid-point of the elastic lithosphere and reach maximal values at its upper and lower boundaries. However, in reality, brittle failure and ductile flow limit the maximum possible stresses at the top and bottom of the mechanical lithosphere, respectively. Although the bottom of the mechanical lithosphere is viscous, it deforms only slowly and thus can also support surface topography or subsurface loads.

Specifically, the mechanical lithosphere encompasses both elastically and viscously strong parts, as defined by a yield-stress envelope. For a given rheological model of the lithosphere, the elastic thickness can be used to determine the mechanical thickness, which in turn is tied to the lithosphere's thermal state (e.g., McNutt 1984). The mechanical lithosphere is always thicker than the elastic lithosphere—but the total bending moments of both
<figure><img src="image_14.png" /><figcaption>**Fig. 12** - Volcanic features on Venus (here, a pancake dome) impose loads on the lithosphere. Even in the absence of gravity data, the topographic response of the lithosphere to these loads places constraints on models of lithospheric thickness. Rheological models enable conversions between inferred elastic and mechanical thicknesses—and then from mechanical thickness to the vertical thermal gradient and thus heat flow. Modified from Borrelli et al. (2021) with illustration by Joanna Wendel</figcaption></figure>

(i.e., integrated over their yield strength envelopes) are equal. As shown in Fig. 12, the bottom of the mechanical lithosphere is associated with the temperature above which the ductile strength drops below a critical value (e.g., ~50 MPa). The corresponding critical temperature depends on the assumed rheology and strain rate but is typically ~1000 K for models with rheologies of dry olivine or diabase (e.g., Molnar, 2020). Stresses in the viscous layer can relax with time, such that the thickness of the mechanical layer can decrease with time. The thickness of the mechanical lithosphere can be used to estimate the thermal gradient due to the strong dependence of rock strength on temperature. Finally, the surface heat flow equals the product of the thermal gradient and the thermal conductivity of the lithosphere. Past papers have used a range of thermal conductivity values, from ~2–4 W m⁻¹ K⁻¹ (e.g., Brown and Grimm 1996; Johnson and Sandwell 1994; Sandwell and Schubert 1992b; Solomon and Head 1982; Reese et al. 1998; O'Rourke and Smrekar 2018; Russell and Johnson 2021; Borrelli et al., 2021; Smrekar et al. 2023). Thus, comparing heat flow values from different studies requires normalizing them to a given thermal conductivity. In this paper, any reference to a heat flow value is normalized to a value of 3 W m⁻¹ K⁻¹.

Strain rate is an essential parameter for calculating the thickness of the mechanical lithosphere and thus thermal gradient and heat flow. Similarly, strain rate directly influences calculated yield stress values used to approximate the complex deformation of the lithosphere in models of convective regime, as it determines whether an active or stagnant lid is predicted. At present there are no direct measurements of strain rates for Venus. Currently the only potential method of accurately constraining local strain rates under consideration for Venus is repeat-pass interferometry (RPI), which could measure active deformation on the timescale of years (Hensley et al., 2022). The strain rate for terrestrial intraplate deformation of ~10⁻¹⁶ s⁻¹ is often applied to Venus as a best guess (e.g. Johnson and Sandwell 1994). However, numerous Earth studies have shown that, in actively deforming regions, and especially volcanically active regions, strain rates can be orders of magnitude higher (e.g., Fagereng and Biggs, 2019; Molnar, 2020) and these locally higher rates should also be
<figure><img src="image_14.png" /><figcaption>**Fig. 12** - Volcanic features on Venus (here, a pancake dome) impose loads on the lithosphere. Even in the absence of gravity data, the topographic response of the lithosphere to these loads places constraints on models of lithospheric thickness. Rheological models enable conversions between inferred elastic and mechanical thicknesses—and then from mechanical thickness to the vertical thermal gradient and thus heat flow. Modified from Borrelli et al. (2021) with illustration by Joanna Wendel</figcaption></figure>

(i.e., integrated over their yield strength envelopes) are equal. As shown in Fig. 12, the bottom of the mechanical lithosphere is associated with the temperature above which the ductile strength drops below a critical value (e.g., ~50 MPa). The corresponding critical temperature depends on the assumed rheology and strain rate but is typically ~1000 K for models with rheologies of dry olivine or diabase (e.g., Molnar, 2020). Stresses in the viscous layer can relax with time, such that the thickness of the mechanical layer can decrease with time. The thickness of the mechanical lithosphere can be used to estimate the thermal gradient due to the strong dependence of rock strength on temperature. Finally, the surface heat flow equals the product of the thermal gradient and the thermal conductivity of the lithosphere. Past papers have used a range of thermal conductivity values, from ~2–4 W m⁻¹ K⁻¹ (e.g., Brown and Grimm 1996; Johnson and Sandwell 1994; Sandwell and Schubert 1992b; Solomon and Head 1982; Reese et al. 1998; O'Rourke and Smrekar 2018; Russell and Johnson 2021; Borrelli et al., 2021; Smrekar et al. 2023). Thus, comparing heat flow values from different studies requires normalizing them to a given thermal conductivity. In this paper, any reference to a heat flow value is normalized to a value of 3 W m⁻¹ K⁻¹.

Strain rate is an essential parameter for calculating the thickness of the mechanical lithosphere and thus thermal gradient and heat flow. Similarly, strain rate directly influences calculated yield stress values used to approximate the complex deformation of the lithosphere in models of convective regime, as it determines whether an active or stagnant lid is predicted. At present there are no direct measurements of strain rates for Venus. Currently the only potential method of accurately constraining local strain rates under consideration for Venus is repeat-pass interferometry (RPI), which could measure active deformation on the timescale of years (Hensley et al., 2022). The strain rate for terrestrial intraplate deformation of ~10⁻¹⁶ s⁻¹ is often applied to Venus as a best guess (e.g. Johnson and Sandwell 1994). However, numerous Earth studies have shown that, in actively deforming regions, and especially volcanically active regions, strain rates can be orders of magnitude higher (e.g., Fagereng and Biggs, 2019; Molnar, 2020) and these locally higher rates should also be
considered for the Venus lithosphere. The implication is that the strength of the lithosphere is likely higher during deformation than previously recognized. Ultimately, based on the effect of strain rate, lithospheric yield strength may be frequently underestimated.

Models of topographic profiles exhibiting flexural bending have been applied to deformation at a variety of features on Venus, including volcanic domes, coronae, and ridge belts. Interestingly, no flexural signatures have been found at larger volcanoes, possibly due to volcanic flooding of the moats (e.g., McGovern and Solomon, 1997). Most coronae have an elastic thickness of 20 km or less, with a handful of greater values mostly at larger coronae (Johnson and Sandwell 1994; Schubert and Sandwell 1995; O'Rourke and Smrekar 2018). Borrelli et al. (2021) examined topographic flexure at steep-sided volcanic domes (diameters of ~20–40 km). Topographic profiles at 20% of the ~60 domes on Venus show evidence of flexure, with most values of elastic thickness in the range from ~15–40 km. For some domes and small-scale features, topographic resolution may be a factor. McGovern et al. (2013) used magma-ascent models to predict that volcanoes may tend to form on thicker lithosphere than typical coronae. The lithospheric thicknesses estimated by Borrelli et al. (2021) generally support this hypothesis. Russell and Johnson (2021) estimated an elastic thickness of ~3–9 km for a small dome that sits on the fracture annulus of Aramaiti Corona. They advocated for the use of a lower Young's modulus and the resulting relatively low elastic thickness due to the presence of major fracturing in the annulus, as well as the best location of flexure per the location of fractures surrounding the dome. Borrelli et al. (2021) also noted that steep-sided domes in the vicinity of coronae tend to have lower elastic thickness than domes elsewhere.

Collectively, these studies found a wide range of heat flow values. The estimates of heat flow associated with coronae are typically ~>50–75 mW m⁻², with a range of ~15 to >200 mW m⁻² (Johnson and Sandwell 1994; Schubert and Sandwell 1995; O'Rourke and Smrekar 2018; Smrekar et al. 2023). Steep-sided domes typically have a lower heat flow than common values for coronae (e.g., ~50 mW m⁻² for domes that are not near coronae; see Borrelli et al., 2021; Russell and Johnson, 2021), unless they are themselves near coronae. Additionally, most of the coronae identified as likely subduction zones (Schubert and Sandwell 1995) generally have greater elastic thicknesses and lower heat flows than average coronae values. Finally, the modelling of tectonic and impact features can also provide estimates of thermal gradient and heat flow. In a recent example, Bjornes et al. (2021) simulated the multi-ring structure of Mead crater and found that the spacing of the ring faults can be explained by a thermal gradient of up to ~14 mK m⁻¹ (or a heat flow of ~42 mW m⁻², for a conductivity of 3 W m⁻¹ K⁻¹), indicating a locally thick lithosphere at the time of formation. In contrast, models of tectonic deformation at other locations point to a very thin lithosphere and high heat flow (e.g., Smrekar and Solomon 1992; Raitala et al. 1995; Ruiz, 2007). As on Earth, the lithospheric thickness and heat flow are likely variable in time and space on Venus.

Elastic lithospheric thickness can also be estimated from the ratio of gravity and topography, or the admittance, due to the wavelength dependence of the deformation response. This method has been applied at numerous highland regions, typically returning estimates of elastic thickness of ~20–30 km (e.g., Smrekar 1994; Simons et al., 1997)—hence the assumption of ~20 km on average by James et al. (2013). Global admittance maps (Anderson and Smrekar 2006; Jiménez-Díaz et al. 2015) find a range of elastic thicknesses of ~0–94 km. Jiménez-Díaz et al. (2015) derive an average value of ~30–40 km, although this value is likely a bit high based on their approach of limiting the minimum value to 14 km.

Heat flow estimates are a strong constraint on the models of convective state and thermal evolution of a planetary body (see Rolf et al. 2022, this collection). Models of mantle con-
considered for the Venus lithosphere. The implication is that the strength of the lithosphere is likely higher during deformation than previously recognized. Ultimately, based on the effect of strain rate, lithospheric yield strength may be frequently underestimated.

Models of topographic profiles exhibiting flexural bending have been applied to deformation at a variety of features on Venus, including volcanic domes, coronae, and ridge belts. Interestingly, no flexural signatures have been found at larger volcanoes, possibly due to volcanic flooding of the moats (e.g., McGovern and Solomon, 1997). Most coronae have an elastic thickness of 20 km or less, with a handful of greater values mostly at larger coronae (Johnson and Sandwell 1994; Schubert and Sandwell 1995; O'Rourke and Smrekar 2018). Borrelli et al. (2021) examined topographic flexure at steep-sided volcanic domes (diameters of ~20–40 km). Topographic profiles at 20% of the ~60 domes on Venus show evidence of flexure, with most values of elastic thickness in the range from ~15–40 km. For some domes and small-scale features, topographic resolution may be a factor. McGovern et al. (2013) used magma-ascent models to predict that volcanoes may tend to form on thicker lithosphere than typical coronae. The lithospheric thicknesses estimated by Borrelli et al. (2021) generally support this hypothesis. Russell and Johnson (2021) estimated an elastic thickness of ~3–9 km for a small dome that sits on the fracture annulus of Aramaiti Corona. They advocated for the use of a lower Young's modulus and the resulting relatively low elastic thickness due to the presence of major fracturing in the annulus, as well as the best location of flexure per the location of fractures surrounding the dome. Borrelli et al. (2021) also noted that steep-sided domes in the vicinity of coronae tend to have lower elastic thickness than domes elsewhere.

Collectively, these studies found a wide range of heat flow values. The estimates of heat flow associated with coronae are typically ~>50–75 mW m⁻², with a range of ~15 to >200 mW m⁻² (Johnson and Sandwell 1994; Schubert and Sandwell 1995; O'Rourke and Smrekar 2018; Smrekar et al. 2023). Steep-sided domes typically have a lower heat flow than common values for coronae (e.g., ~50 mW m⁻² for domes that are not near coronae; see Borrelli et al., 2021; Russell and Johnson, 2021), unless they are themselves near coronae. Additionally, most of the coronae identified as likely subduction zones (Schubert and Sandwell 1995) generally have greater elastic thicknesses and lower heat flows than average coronae values. Finally, the modelling of tectonic and impact features can also provide estimates of thermal gradient and heat flow. In a recent example, Bjornes et al. (2021) simulated the multi-ring structure of Mead crater and found that the spacing of the ring faults can be explained by a thermal gradient of up to ~14 mK m⁻¹ (or a heat flow of ~42 mW m⁻², for a conductivity of 3 W m⁻¹ K⁻¹), indicating a locally thick lithosphere at the time of formation. In contrast, models of tectonic deformation at other locations point to a very thin lithosphere and high heat flow (e.g., Smrekar and Solomon 1992; Raitala et al. 1995; Ruiz, 2007). As on Earth, the lithospheric thickness and heat flow are likely variable in time and space on Venus.

Elastic lithospheric thickness can also be estimated from the ratio of gravity and topography, or the admittance, due to the wavelength dependence of the deformation response. This method has been applied at numerous highland regions, typically returning estimates of elastic thickness of ~20–30 km (e.g., Smrekar 1994; Simons et al., 1997)—hence the assumption of ~20 km on average by James et al. (2013). Global admittance maps (Anderson and Smrekar 2006; Jiménez-Díaz et al. 2015) find a range of elastic thicknesses of ~0–94 km. Jiménez-Díaz et al. (2015) derive an average value of ~30–40 km, although this value is likely a bit high based on their approach of limiting the minimum value to 14 km.

Heat flow estimates are a strong constraint on the models of convective state and thermal evolution of a planetary body (see Rolf et al. 2022, this collection). Models of mantle con-
vection estimate thermal lithospheric thickness using geoid and topography data and magmatic activity as constraints (see Rolf et al. 2022, this collection). Using geoid-topography ratios, the average thermal lithospheric thickness may be 200 to 400 km, with reduced thickness below volcanic highlands (Moore & Schubert, 1997; Solomatov and Moresi, 1996). Larger average thicknesses of up to ~600 km may be feasible if geoid and topography can be explained by the thermal isostasy of a stagnant lid (Orth & Solomatov, 2011). The unknown temperature and viscosity of Venus' mantle allow for a very wide range of values from such models.

In contrast, a thinner thermal lithosphere of ~100–150 km is required to permit pressure-release melting at Venus hot spots (e.g., Smrekar & Parmentier 1996; Nimmo and McKenzie 1998), and to agree with estimates of heat flow (Smrekar et al. 2023). These estimates are compatible with mantle temperatures calculated from inferred compositions of basalts at the surface (e.g., Nimmo, 2002; Lee et al. 2009; Filiberto, 2014). Estimates of the mantle potential temperature ($T_P$) (i.e., the mantle adiabat projected up to the surface) have been derived based on data from Venera 13 ($T_P$ = 1732 ± 73 K) and Venera 14 ($T_P$ = 1603 K, 1643 ± 70 K, 1732 ± 101 K)—within the range of ambient conditions of Earth (i.e., $T_P$ = 1623 ± 50 K). Of note, estimates based on measurements from Vega 2 ($T_P$ = 2051 ± 167 K) may be substantially higher but could be within the range of Earth's ambient mantle, too (e.g., McKenzie et al. 1992; Lee et al. 2009; Weller et al., 2015; Shellnutt 2016).

On Mars, where surface ages span several Gyr and flexural features can be dated back to the first ~1–3 Ga, it is possible to show a history of cooling with time based on increasing lithospheric thickness with the age of the terrain (e.g., McGovern et al., 2004). The inability to date individual areas on Venus means that different lithospheric thickness on Venus globally could reflect spatial or temporal variations. Given Venus' young surface age, the average lithospheric thickness may not have increased significantly over the last ~500–1000 Myr. The question remains as to whether the estimates of elastic thickness and heat flow reflect the present-day lithosphere or the time of lithospheric loading. In areas of thicker lithosphere, the loading signature from the time of loading may be 'frozen in' even if the lithosphere continues to cool and thicken with time. Smrekar et al. (2022b) argue that some regions with thin lithosphere and high heat flow, such Parga Chasma, are likely to represent areas with high heat flow today. Even given some uncertainties, lithospheric thickness and heat flow can inform models of the present-day lithosphere and interior (e.g., Rolf et al. 2022, this collection).

# 5 Models for the Formation of Volcanic and Tectonic Features

Understanding the evolution of Venus and planets in general requires modelling of specific geologic processes to explore plausible conditions in the mantle (potential temperature, de-compression melting including composition and volatile content, etc.) and the lithosphere (mechanical thickness, strength, thermal gradient, stress state, etc.). The key constraints de-scribed above form the recent or present conditions that evolutionary models should match. In this section, we discuss models of the origin of specific features and the conditions im-plied over the resurfacing age of Venus up to the present. We then use all these inputs to ex-plore the hypothesis that conditions on Venus currently are much like those on early Earth. Rolf et al. 2022 (this collection) examines models of Venus' interior evolution, including the deep parts of the planet, across a billion-year timescale.
vection estimate thermal lithospheric thickness using geoid and topography data and magmatic activity as constraints (see Rolf et al. 2022, this collection). Using geoid-topography ratios, the average thermal lithospheric thickness may be 200 to 400 km, with reduced thickness below volcanic highlands (Moore & Schubert, 1997; Solomatov and Moresi, 1996). Larger average thicknesses of up to ~600 km may be feasible if geoid and topography can be explained by the thermal isostasy of a stagnant lid (Orth & Solomatov, 2011). The unknown temperature and viscosity of Venus' mantle allow for a very wide range of values from such models.

In contrast, a thinner thermal lithosphere of ~100–150 km is required to permit pressure-release melting at Venus hot spots (e.g., Smrekar & Parmentier 1996; Nimmo and McKenzie 1998), and to agree with estimates of heat flow (Smrekar et al. 2023). These estimates are compatible with mantle temperatures calculated from inferred compositions of basalts at the surface (e.g., Nimmo, 2002; Lee et al. 2009; Filiberto, 2014). Estimates of the mantle potential temperature ($T_P$) (i.e., the mantle adiabat projected up to the surface) have been derived based on data from Venera 13 ($T_P$ = 1732 ± 73 K) and Venera 14 ($T_P$ = 1603 K, 1643 ± 70 K, 1732 ± 101 K)—within the range of ambient conditions of Earth (i.e., $T_P$ = 1623 ± 50 K). Of note, estimates based on measurements from Vega 2 ($T_P$ = 2051 ± 167 K) may be substantially higher but could be within the range of Earth's ambient mantle, too (e.g., McKenzie et al. 1992; Lee et al. 2009; Weller et al., 2015; Shellnutt 2016).

On Mars, where surface ages span several Gyr and flexural features can be dated back to the first ~1–3 Ga, it is possible to show a history of cooling with time based on increasing lithospheric thickness with the age of the terrain (e.g., McGovern et al., 2004). The inability to date individual areas on Venus means that different lithospheric thickness on Venus globally could reflect spatial or temporal variations. Given Venus' young surface age, the average lithospheric thickness may not have increased significantly over the last ~500–1000 Myr. The question remains as to whether the estimates of elastic thickness and heat flow reflect the present-day lithosphere or the time of lithospheric loading. In areas of thicker lithosphere, the loading signature from the time of loading may be 'frozen in' even if the lithosphere continues to cool and thicken with time. Smrekar et al. (2022b) argue that some regions with thin lithosphere and high heat flow, such Parga Chasma, are likely to represent areas with high heat flow today. Even given some uncertainties, lithospheric thickness and heat flow can inform models of the present-day lithosphere and interior (e.g., Rolf et al. 2022, this collection).

# 5 Models for the Formation of Volcanic and Tectonic Features

Understanding the evolution of Venus and planets in general requires modelling of specific geologic processes to explore plausible conditions in the mantle (potential temperature, de-compression melting including composition and volatile content, etc.) and the lithosphere (mechanical thickness, strength, thermal gradient, stress state, etc.). The key constraints de-scribed above form the recent or present conditions that evolutionary models should match. In this section, we discuss models of the origin of specific features and the conditions im-plied over the resurfacing age of Venus up to the present. We then use all these inputs to ex-plore the hypothesis that conditions on Venus currently are much like those on early Earth. Rolf et al. 2022 (this collection) examines models of Venus' interior evolution, including the deep parts of the planet, across a billion-year timescale.
# 5.1 Formation of Volcanic Features

Despite their broad similarities, large volcanoes on Venus appear to have some differences in their formation relative to their Terran and Martian counterparts (Fig. 1). Higher surface temperatures and pressures on Venus today yield smaller edifice heights than would otherwise be the case (e.g., Head and Wilson 1986; Bridges 1997; Head and Wilson 1992). Higher surface temperatures lead to smaller thermal gradients, shallower magma sources, hotter erupting lava, and therefore lava flows that spread farther laterally than on Earth (e.g., Flynn et al. 2023). The capacity of the thick, dense atmosphere to transport heat by radiation and convection also affects the spreading of lava flows (e.g., Snyder 2002; Flynn et al. 2023). The high surface pressure of Venus was also found to inhibit magma volatile exsolution, leading to fewer gas bubbles and thus potentially to less explosive volcanic eruptions than on Earth (e.g., Airey et al. 2015).

## 5.1.1 Plains Topography and Volcanism

Most of Venus' surface consists of volcanic plains. The plains are defined as regions of low elevation, low relief topography. As described above, there are a wide range of volcanic features in the plains. Studies to date have examined the formation of these plains via investigating the origin of the low lying topography as a possible result of thermal isostasy, and via numerical models that predict either global or regional volcanism. Globally, Venus' hypsometry is largely unimodal, lying within ±1 km of the mean planetary radius (6051.8 km, although see Smrekar et al. 2018 for a discussion of elevations above 2 km). The hypsometry of the lowlands—and thus of the volcanic plains—is well-fit by a straight line (Rosenblatt et al., 1994), indicating that these regions are dominantly due to thermal isostasy. For an Earth-like mantle temperature, a thermal isostasy model applied to Venus gives a thermal lithospheric thickness of ~90 km. Means other than plate tectonic spreading centres can produce thermal isostasy (e.g., Morgan and Phillips 1983; Rosenblatt et al. 1994). For example, Morgan and Phillips (1983) showed that hot spot volcanism could produce the hypsometry of the plains. Ghail (2015) explained the plains topography as lithospheric rejuvenation—proposing that Venus may have a thin crust underlain by a localized asthenosphere activated by the presence of CO₂ in regions where the heat flow is relatively high, such as at mantle upwellings, plumes, or hot spots. Local topographic deviations from the global pattern of up to of several 100 m, such as those in Baltis Vallis, may be due to current convection-driven dynamic compensation of the topography (e.g. Conrad and Nimmo 2023).

Numerous geodynamic models predict widespread volcanism, or plains volcanism, and resurfacing, as discussed in detail in Rolf et al. 2022. These models predict volcanism and crustal thickness as a function of the lithospheric thickness, mantle temperature, and volatile content. We describe some example models here for completeness. In terms of explaining the occurrence of global plain volcanism on Venus, Parmentier and Hess (1992) and Head et al. (1994) presented a model linking the vertical accretion of a basaltic crust and the subsequent delamination linked to the emplacement of widespread volcanic plains. A 'catastrophic' tectonic resurfacing model initiated by delamination of negatively buoyant, depleted mantle was proposed to cause upwelling of warm, fertile mantle and pressure-release melting to produce extensive surface volcanism. Recent geodynamic modelling work by Adams et al. (2022) also focused on lithospheric delamination, in a 'peel-back' tectonic style, to be followed by the emplacement of hot, buoyant asthenosphere beneath the crust that may give rise to regional-scale volcanism. This style of resurfacing is proposed to occur on a more regional scale rather than a 'catastrophic' one.
# 5.1 Formation of Volcanic Features

Despite their broad similarities, large volcanoes on Venus appear to have some differences in their formation relative to their Terran and Martian counterparts (Fig. 1). Higher surface temperatures and pressures on Venus today yield smaller edifice heights than would otherwise be the case (e.g., Head and Wilson 1986; Bridges 1997; Head and Wilson 1992). Higher surface temperatures lead to smaller thermal gradients, shallower magma sources, hotter erupting lava, and therefore lava flows that spread farther laterally than on Earth (e.g., Flynn et al. 2023). The capacity of the thick, dense atmosphere to transport heat by radiation and convection also affects the spreading of lava flows (e.g., Snyder 2002; Flynn et al. 2023). The high surface pressure of Venus was also found to inhibit magma volatile exsolution, leading to fewer gas bubbles and thus potentially to less explosive volcanic eruptions than on Earth (e.g., Airey et al. 2015).

## 5.1.1 Plains Topography and Volcanism

Most of Venus' surface consists of volcanic plains. The plains are defined as regions of low elevation, low relief topography. As described above, there are a wide range of volcanic features in the plains. Studies to date have examined the formation of these plains via investigating the origin of the low lying topography as a possible result of thermal isostasy, and via numerical models that predict either global or regional volcanism. Globally, Venus' hypsometry is largely unimodal, lying within ±1 km of the mean planetary radius (6051.8 km, although see Smrekar et al. 2018 for a discussion of elevations above 2 km). The hypsometry of the lowlands—and thus of the volcanic plains—is well-fit by a straight line (Rosenblatt et al., 1994), indicating that these regions are dominantly due to thermal isostasy. For an Earth-like mantle temperature, a thermal isostasy model applied to Venus gives a thermal lithospheric thickness of ~90 km. Means other than plate tectonic spreading centres can produce thermal isostasy (e.g., Morgan and Phillips 1983; Rosenblatt et al. 1994). For example, Morgan and Phillips (1983) showed that hot spot volcanism could produce the hypsometry of the plains. Ghail (2015) explained the plains topography as lithospheric rejuvenation—proposing that Venus may have a thin crust underlain by a localized asthenosphere activated by the presence of CO₂ in regions where the heat flow is relatively high, such as at mantle upwellings, plumes, or hot spots. Local topographic deviations from the global pattern of up to of several 100 m, such as those in Baltis Vallis, may be due to current convection-driven dynamic compensation of the topography (e.g. Conrad and Nimmo 2023).

Numerous geodynamic models predict widespread volcanism, or plains volcanism, and resurfacing, as discussed in detail in Rolf et al. 2022. These models predict volcanism and crustal thickness as a function of the lithospheric thickness, mantle temperature, and volatile content. We describe some example models here for completeness. In terms of explaining the occurrence of global plain volcanism on Venus, Parmentier and Hess (1992) and Head et al. (1994) presented a model linking the vertical accretion of a basaltic crust and the subsequent delamination linked to the emplacement of widespread volcanic plains. A 'catastrophic' tectonic resurfacing model initiated by delamination of negatively buoyant, depleted mantle was proposed to cause upwelling of warm, fertile mantle and pressure-release melting to produce extensive surface volcanism. Recent geodynamic modelling work by Adams et al. (2022) also focused on lithospheric delamination, in a 'peel-back' tectonic style, to be followed by the emplacement of hot, buoyant asthenosphere beneath the crust that may give rise to regional-scale volcanism. This style of resurfacing is proposed to occur on a more regional scale rather than a 'catastrophic' one.