PHYSICAL CONTROLS ON MAGMAS IN MONOGENETIC VOLCANIC SYSTEMS: FROM MAGMA GENERATION TO SURFACE ERUPTION

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Part 2 Physical controls on magmas in monogenetic volcanic systems: from magma generation to surface eruption.

Introduction

This review summarizes the current state of knowledge on the physical controls in a monogenetic volcanic system, which include: (1) the melting anomaly at the magma source; (2) the conduits through which magma rises to the surface; (3) the magma chambers (if any are formed); (4) the zones of geothermal influence; and (5) the volcanic edifice [Walker, 1993; 2000]. The generation, segregation, and propagation of mafic magmas, from the mantle to the surface of the Earth, involve different time and spatial scales [Walker, 1993; 2000; Canon-Tapia and Walker, 2004]. An understanding of the physical controls on these processes is fundamental to understanding the occurrence and development of monogenetic basaltic volcanism anywhere around the world. Mafic magmas are generated by partial melting in the upper mantle [O’Hara, 1965]. The first melts are formed in a partially molten mantle matrix and are considered to have an ascent rate on the order of only several to tens of centimetres per year due to the high viscosities of the mantle matrix involved [Rubin, 1993a; Hewitt, 2010]. Melt segregation is also a slow process occurring on the order of millimetres per year [Faul, 1997]. On the other hand, basaltic volcanic eruptions liberate large volumes of mafic magmas and last from days to years, with dikes propagating on the order of meters per second [Battaglia et al., 2005; Demouchy et al., 2006]. Fast extraction of large amounts of mafic magma from the mantle is thus considered to be controlled by hydraulic fracturing driven by magma within cracks (dikes) [e.g., Rubin, 1995; and references therein]. The various processes involved in a volcanic system are listed in Table 2-1. The reasons magma accumulates and then propagates toward the surface, as well as the nature of the principal drivers of its propagation, are still debated [Valentine, 1989], despite the importance of these key processes for predicting future eruptions in monogenetic basaltic fields. In addition to the controls on magma propagation at depth,
another concern arises when one looks at the development of a monogenetic basaltic field at the Earth’s surface. This type of volcanism is characterized by small volcanic centres (i.e., scoria cones, maars, tuff rings and shield volcanoes) scattered across a restricted area (i.e., a field). In general, each volcano in a monogenetic basaltic volcanic field is related to one single eruption [Brenna et al., 2011], corresponding to a single transfer of magma from a source at depth up to the surface. Therefore, an important question is what causes each magma batch to extrude at a different location. To set the framework for the thesis, this Part describes the general physical controls acting on the magma from its source to surface. Firstly, the rheological characteristics of the medium hosting the magma from source to surface are summarized, and then different theories on how and where magmas are generated are discussed. Finally, theories on how magmas segregate and propagate are presented. This review highlights recent research into the controls on the ascent of magmas, and touches on important gaps in our understanding of monogenetic basaltic volcanism, which will be addressed in subsequent Parts.

Rheology of the lithosphere

A monogenetic basaltic volcanic system comprises the generation of magmas in the mantle and the propagation of these magmas through each constituent layer of the lithosphere (upper mantle, lower and upper crust), which influences the viability of the system. The generation, segregation, and propagation of magmas are controlled by the rheological properties of all the lithosphere layers.

Theory

‘Rheology’ describes the response of materials, such as rocks or melts, to an imposed stress regime. The response depends on the intrinsic properties of the material (e.g., composition, density) and extrinsic conditions (e.g., pressure, temperature, chemistry) [Ranalli, 1995]. Materials can be homogeneous (with isotropic or anisotropic mechanical properties) or inhomogeneous, and are characterized by parameters such as elasticity, rigidity, compressibility, and viscosity [Hawkes and Mellor, 1970; McBirney and Murase, 1984; Ranalli, 1995; Ancey, 2007; Burov, 2011]. Three dominant rheological behaviours, elasticity, viscous flow and plasticity, ideally characterize relationships between stress, strain and time (Fig. 2-1). Most materials can support small strains without being permanently deformed (elastic behaviour). However, if the stress exceeds their yield strength (threshold at which permanent deformation occurs), either brittle rupture (for purely elastic materials), or plastic flow will occur, depending on the environment and the composition of the material (Fig. 2-2a). The brittle mode occurs at low confining pressure and low temperature and/or high differential stress and/or high strain rates. Plastic flow is ductile, and occurs in high temperature environments and/or low differential stress and/or low strain rates. Plastic deformation includes a number of different creep2 mechanisms, depending on the differential stress and temperature (Fig. 2-2a):
i) diffusion creep (volume diffusion (Nabarro-Herring creep), grain boundary diffusion (Cobble creep), pressure solution and grain boundary sliding,
ii) dislocation glide and
iii) dislocation creep (dislocation climb and dislocation cross-slip) [Karato, 2008; Fossen, 2010]. Brittle and ductile environments are separated by the brittle-ductile transition (BDT) zone. The BDT zone depends upon pressure, temperature, strain rate and composition of the constituent rocks.

Strength of the lithosphere

The lithosphere is the strong outer rind of the solid Earth, which exerts an important influence on geodynamic environments. The strength of the lithosphere represents, in a simple model, the critical yield strength permitting a plate tectonic style of convection [Karato, 2008]. The strength of the lithosphere is defined by the strength of each constituent layer, predominantly controlled by a single mineral: quartz in the upper crust, feldspar in the lower crust and olivine in the mantle [Dobran, 2001]. The type of deformation affecting the constituent minerals is a function of temperature, pressure, differential stress, strain rate and hydrostatic pressure. Stress – strain – temperature data for individual minerals can be displayed in rock deformation maps (Fig. 2-2b, c). The differential stress on rocks as a function of temperature for different strain rates delineates different regions of creep present in the material under different loading conditions [Davis and Reynolds, 1996]. From these behaviours, the conventional models for the strength of the lithosphere are visualized using the Brace-Goetze strength profile, or “Christmas Tree” profile [Ranalli and Murphy, 1987] (Fig. 2-3). These profiles combine the pressure-dependent Byerlee’s  friction law, and the temperature- and strain-dependent creep mechanisms [Ranalli and Murphy, 1987]. Although widely accepted as representative of the rheological behaviour of the lithosphere, these “Christmas tree” profiles only show a quasi-static view of the rheological response to stress, and do not take into account the dynamic interactions between the brittle and ductile behaviours of the constituent layers of the lithosphere [Regenauer-Lieb et al., 2006], neither the influence of the strain rate [Lubarda et al., 2003]. In addition, the assumption of monomineralic rocks is also a limitation: natural rocks have generally a second crystalline phase with different rheological behaviour with might influence the shape of the profile [Rutter and Brodie, 1991].

 Properties of the lithospheric mantle

The principal components of the lithospheric mantle are olivine, pyroxene and garnet; because olivine is weaker than other phases and volumetrically dominant, the deformation of the upper mantle is assumed to be rheologically controlled by olivine [Karato and Wu, 1993; Karato, 2008; Karato et al., 2008; Demouchy et al., 2009]. The mantle exhibits different mechanical properties depending on the time scale of mantle deformation processes. The mantle has linear (Newtonian) viscosity for times longer than about 103 years corresponding to the Maxwell relaxation time (Fig. 2-4). Diffusion creep and dislocation creep are the two most important flow mechanisms. For fine
grained rocks and low stress, diffusion creep is dominant and the mantle behaves as a Newtonian fluid. For high stress, however, dislocation creep dominates, and the mantle behaves as a power-law fluid (Fig. 2-1; Karato and Wu, 1993). Diffusion creep results in no or weak lattice-preferred orientation4 (LPO), whereas dislocation creep results in strong LPO, which is an important mechanism in melt formation and connectivity (see section 2.4.1) [Kohlstedt and Holtzman, 2009].
Lithospheric mantle behaviour follows viscoelastic rheological models (Fig. 2-1) [Karato and Wu, 1993; Burgmann and Dresen, 2008; Karato et al., 2008; Karato, 2010]. If transient viscous deformation is neglected (e.g., viscous flow following earthquakes, Decriem and Árnadóttir, in press), then the mantle can be considered to show short term elastic response followed by long-term linearly viscous response in the form of viscoelastic behaviour, and can be compared to a Maxwell body type (Fig. 2-1) [Burgmann and Dresen, 2008]. A transient response can be introduced by adding a Kelvin element in series with the Maxwell element, forming a Burger’s body (Fig. 2-1). If the mantle flow is considered to be linear, then the Burger’s body is the most complete
and simplest rheological model for its behaviour (taking into account the elastic, transient and steady-state deformation). However, despite being a good representation,

Table of Content
ABSTRACT
ACKNOWLEDGMENTS
TABLE OF CONTENT
PART 1 INTRODUCTION
1.1 Background
1.2 Thesis outline
PART 2 PHYSICAL CONTROLS ON MAGMAS IN MONOGENETIC VOLCANIC SYSTEMS: FROM MAGMA GENERATION TO SURFACE ERUPTION
2.1 Introduction
2.2 Rheology of the lithosphere
2.3 Fracturing
2.4 Ascent history of a magma batch
2.5 Discussion
PART 3 SPATIAL DISTRIBUTION AND ALIGNMENTS OF VOLCANIC CENTERS: CLUES ON THE FORMATION OF MONOGENETIC VOLCANIC FIELDS.
3.1 Introduction
3.2 Database and methods
3.3 Limitations of the analytical methods
3.4 Results
3.5 Discussion
3.6 Conclusions
PART 4 INTERACTION OF ASCENDING MAGMA WITH PRE-EXISTING CRUSTAL  FRACTURES IN MONOGENETIC BASALTIC VOLCANISM: AN EXPERIMENTAL  APPROACH 
4.1 Introduction
4.2 Analogue modeling
4.3 Stress regimes
4.4 Limitations
4.5 Observations and results
4.6 Implications for the interaction of propagating dikes and pre-existing fractures
4.7 Conclusion
PART 5 AGE, DISTANCE AND GEOCHEMISTRY: ANALYSING PATTERNS IN THE AUCKLAND VOLCANIC FIELD ERUPTION SEQUENCE
5.1 Introduction
5.2 Database
5.3 Methodology
5.4 Results
5.5 Discussion
5.6 Conclusion
PART 6 CONCLUSIONS
6.1 Outlook
REFERENCES

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Physical and structural controls on monogenetic basaltic volcanism, and implications for the evolution of the Auckland Volcanic Field

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