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Relations between deformation and upper crustal magma emplacement in laboratory physical models
Tectonophysics 484 (2010) 139–146
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Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o
Relations between deformation and upper crustal magma emplacement in laboratory physical models F. Mazzarini a,⁎, G. Musumeci b, D. Montanari c, G. Corti d a
Istituto Nazionale di Geofisica e Vulcanologia, Via della Faggiola 32, 56126 Pisa, Italy Dipartimento di Scienze della Terra, Università di Pisa, Via S. Maria 54, 56126, Pisa, Italy Centro di Eccellenza per la Geotermia di Larderello, Via G. Carducci, 4, 56044, Larderello (PI), Italy d Istituto di Geoscienze e Georisorse, CNR, Via G. la Pira 4, 50121 Firenze, Italy b c
a r t i c l e
i n f o
Article history: Received 19 January 2009 Received in revised form 10 June 2009 Accepted 15 September 2009 Available online 23 September 2009 Keywords: Mechanical layering of upper crust Magma emplacement Analogue modelling
a b s t r a c t This paper presents analogue models for the emplacement of granitic magmas in upper crustal levels with different mechanical layering during shortening, extension and strike–slip deformation. In particular, we investigated how a weak layer embedded in the upper brittle crust can control the level of magma emplacement. The adopted experimental setup was used to examine the control of soft rocks on the movement of magma through a deforming brittle crust. Model results indicate that the occurrence of a weak (soft) layer embedded in brittle (stiff) material has an impact on the level of magma emplacement. The level of emplacement during both extension and shortening was systematically deeper for models with a soft layer than for purely brittle models. During strike–slip deformation the magma pierced the surface in both purely brittle and brittle–ductile models. © 2009 Elsevier B.V. All rights reserved.
1. Introduction Magmatism occurs in different geodynamic settings, producing magmas that are emplaced within the crust at levels that may vary from the middle–lower crust to the surface (volcanism). In the continental crust, magma emplacement is well documented and modelled for crustal shortening (Hutton, 1997; Kalakay et al., 2001; Musumeci et al., 2005; Tibaldi, 2005; Galland et al., 2007a,b; Mazzarini et al., 2008), crustal extension (Roma'n-Berdiel, 1999; Corti et al., 2003) and wrenching (e.g. D'Lemos et al., 1992; Hutton and Reavy, 1992; Brown, 1994; Vigneresse, 1995; Roma´n-Berdiel et al., 1997; Salvini et al., 1997; Rosenberg, 2004; Corti et al., 2005). These examples document the even occurrence of intrusions (at depth in the crust) and volcanoes (at the surface) in quite different tectonic settings (convergent, divergent and strike–slip), suggesting that the crustal stress regime does not exert a first order control on the emplacement level of magma within the crust, whereas local strain distribution and near-field stress could provide room for magma emplacement. It is more likely that the level of magma emplacement is controlled by the magma supply rate and crustal rheology at least for continuous magma supply (e.g. Petford et al., 2000). This is particularly evident in the middle and upper crust, where crustal heterogeneities such as
⁎ Corresponding author. Tel.: +39 0508311956; fax +39 0508311942. E-mail address: [email protected] (F. Mazzarini). 0040-1951/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2009.09.013
lithological and rheological variations lead to mechanical weaknesses, which largely determine the depth of magma emplacement (e.g. Roma'nBerdiel et al., 1995, 1997; Kavanagh et al., 2006). The role of mechanical weaknesses seems to be more important in the upper crust, where the formation of sills often occurs by magma inflation and roof uplift, exploiting lithological variations. This is testified by the occurrence of several intrusive bodies at shallow depths within brittle–ductile rheological systems such as sedimentary covers, which contain strength anisotropies that stop the ascent of granitic magmas, allowing them to spread horizontally (Roma'n-Berdiel et al., 1995, 1997; Kavanagh et al., 2006) as well as for basaltic magmas (Pasquarè and Tibaldi, 2007; Tibaldi et al., 2008; Tibaldi and Pasquarè, 2008). The above-documented close spatial relationship between the mechanical layering of the crust and the level of magma emplacement is such that the mechanism which allows the migration of magma through a layered crust deserves further investigation. We therefore investigated the role of mechanical discontinuities in the upper brittle crust during the emplacement of magma in compressive, extensional and strike–slip regimes. To this end, analogue experiments were performed to simulate magma emplacement within a deforming brittle crust with and without a soft layer. The experimental apparatus as well as the crust and magma analogues are subsequently described. Lastly, the results of the analogue experiment are discussed, focussing on the level of emplacement and on the observed relationships between the simulated intrusions and the structures formed in the simulated deformed brittle crust.
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2. Analogue modelling We tested a series of small-scale models to investigate the influence of the mechanical layering of the upper brittle crust on the level of magma emplacement during deformation. The experiments were performed at the Tectonic Modelling Laboratory of the Institute of Geosciences and Earth Resources (National Research Council of Italy) and of the Florence University Department of Earth Sciences. Models, with initial dimensions of 60 cm × 45 cm and a thickness of 5 cm, were built above the metal base plate of the apparatus (Fig. 1). They were designed to reproduce magma intrusion both in a purely brittle crust and in an upper brittle crust comprising weak ductile layers. In the experiments, the brittle behaviour was reproduced using a dry quartzsand characterized by a grain size b250 µm, an internal friction angle of 39°, cohesion of 65 Pa and density of 1550 kg m− 3. The weak layers were reproduced using polydimethilsyloxane (PDMS), a transparent silicone with a density of 965 kg m− 3, a viscosity of 3 × 104 Pas, and Newtonian behaviour at strain rates b3 × 10− 2 s− 1 (Weijermars, 1986). A low viscosity mixture of silicone (Wacker Silicone Bouncing Putty 29) and oleic acid (4:1% in weight) was employed to simulate granitic magma; this mixture exhibits Newtonian behaviour (with viscosity of 7 × 102 Pa s) and a density of 1060 kg m− 3. We performed three different experimental series representing three different tectonic regimes (strike–slip, shortening and extension); for each tectonic regime we tested two models constructed with and without the embedded ductile layer. In the case of shortening, we also tested a
model in which one half has a purely brittle rheology and the other has an embedded ductile layer. Each half has an injection point located at the centre of the base of the model. A moving wall driven by a stepper motor controlled by a central unit enabled deformation of the models (Fig. 1a). Magma intrusion during deformation was simulated using a special injection apparatus consisting of a piston (Fig. 1b) and a magma distribution system made of pipes with a central injection point (10 mm in diameter) on the metal base plate of the apparatus (Fig. 1c). The deformation rate (2 cm/ h) and the magma injection velocity (20 cm/h) remained constant in all the experiments. 2.1. Scaling The experiments were scaled taking into consideration the conditions of geometric, dynamic, kinematic and rheologic similarity outlined by Hubbert (1937), Ramberg (1981) and Weijermars and Schmeling (1986). Geometric similarity was achieved by imposing a scaling ratio of 10− 5, such that 1 cm in the experiments corresponds to 1 km in nature. Density of modelling materials imposes a density scale ratio of ∼0.4–0.5; together with the reduction factor of density and gravity (=1), the resulting model to nature ratio of stresses was of the order of 4 × 10− 6. The scaling ratio of time is ∼1.5 × 10− 9 (1 h in the models represents ∼ 80,000 yrs in nature) and the resulting scaling factor of the linear injection velocity is ∼6.6 × 103. The scaling ratio of viscosity (∼ 6 × 10− 15) results in a scaled viscosity of ∼5 × 1018 Pa s for
Fig. 1. Details of the experimental set up. (a) Oblique view of the deformation apparatus; (b) analogue magma injection piston; (c) metal base of the apparatus showing the magma injection point and the moving wall/fixed base (note that the moving plate attached to the moving wall is used in strike–slip experiments only); (d) schematic model cross-section showing the rheological stratification of models. In the experiments, the modelling length ratio is l⁎ = 1 × 10− 5 (1 cm in the model represents about 1 km in nature).
F. Mazzarini et al. / Tectonophysics 484 (2010) 139–146 Table 1 Main scaling parameters (see Corti et al., 2005). Parameter
Model
Nature
Model / Nature
Length (m) Viscosity (ductile layer) (Pa s) Viscosity (magma) (Pa s) Displacement velocity Emplacement velocity
0.01 3000 700 2 (cm h− 1) 20 (cm h− 1)
1000 ∼5 × 1018 Pa s ∼1017 Pa s 2.7 (cm yr− 1) ∼30 (cm yr− 1)
1 × 10− 5 ∼6 × 10− 15 ∼6 × 10− 15 ∼6.6 × 103 ∼6.6 × 103
the weak ductile layer and ∼ 1017 Pa s for the granitic magma. Scaling parameters for viscosity and for the magma input rate and deformation are reported in Table 1. Details of the scaling procedure are reported in Montanari et al. (2010-this volume). 2.2. Limitations of modelling The performed analogue experiments are only an approximation of nature, and the following should be borne in mind: (i) only simple, purely convergent, divergent or strike–slip tectonic settings were investigated; (ii) the scaling is based on mechanical principles, not thermal ones. In our models, we assume isothermal conditions and do not take into account the thermal effects of intrusion-related heating and pluton cooling on the rheology of the crust and magma, respectively. This assumption is based on the fact that the magma-related thermal perturbation (contact aureole) affects small portions of the host rocks and that magmas may not freeze quickly in actively deforming regions, owing to persisting thermal anomalies related to the repeated intrusion of magmas into the crust. We are aware that, as consequence of this assumption, in our experiment the deformation in the analogue crust is essentially accomplished by faulting and bending without formation of cleavage and/or ductile deformation because we are not able to reproduce the formation of a thermal aureole that could accommodate the high strain induced in the host rock by the magma emplacement (e.g. Albertz et al., 2005); (iii) although the low viscosity mixture of silicone and oleic acid reproduces a significant viscosity contrast between the low viscosity melt fraction and the crust, when scaled to nature, it is only able to simulate a high viscosity magma. This implies that the analogue magma does not move upward through hydrofracturing alone, but also in diapirs. A section of an experiment performed in compression on a purely brittle modelled crust is shown in Fig. 2. The analogue magma rises along the main thrust and locates in the core of the push-up structure where local strain creates room for magma. The adopted material and scaling thus determine magma movement in our experiment by pervasive flow through shear zones and diapirims. Nonetheless, we consider that in these experiments the control of
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mechanical layering on the magma emplacement is always accounted for; (iv) as each layer in the experimental setup is homogeneous, the role of inherited tectonic structures is neglected. Notwithstanding the high mechanical heterogeneity of the upper crust, this assumption is acceptable because we mainly aimed to investigate the level of magma emplacement as a function of interplay between mechanical layering and tectonic regime. (v) magma emplacement in a deforming crust is the result of the interaction of several processes (magma emplacement and cooling, metamorphism and deformation) having different time scales and rates (e.g. Paterson and Tobish, 1992). The injection rates (i.e. the rate at which the magma is introduced in the system), once properly scaled, are in the 20–106 cm yr− 1 range. The scaled deformation rate used in the experiments, is 2.7 cm yr− 1, and the resulting scaled magma supply is in the range 10− 2–10− 4 km3 yr− 1. These values are compatible with rates of faulting, magma supply and magma emplacement observed in nature (Paterson and Tobish, 1992 and references therein; Jellinek and DePaolo, 2003). (vi) the rate at which pluton grows is a critical point. Large batholiths, for example the Tuolumme Intrusive Suite in California, grew by incremental assemblage of small magma batches over millions years (Coleman et al., 2004; Glazner et al., 2004). On the other hands it has been documented that ascent of magma trough the crust and magma emplacement could be a very fast processes (e.g. Petford et al, 2000). In our experiment we consider continuous magma supply of relatively small magma batches (low magma supply, see point above) at very shallow crustal level. 3. Experimental results For each experiment, deformation was monitored through direct observation and sequential top-view photos. At the end of each experiment, the model was sectioned and the level of magma emplacement (height of the magma above the base of the model) was measured. As magma reaches the model surface the height of the surface has been used as the level of magma emplacement. In all the experiments the lowermost boundary of analogue intrusions is at the base of the model, that is the level where the injection point is located. 3.1. Intrusion shape For each deformation setting, it is well known that the shape of the intrusion is primarily controlled by the ratio between the intrusion and deformation velocity (e.g. Roma'n-Berdiel et al., 1997; Corti et al., 2005; Montanari et al., 2010-this volume). An increase in this ratio (i.e., predominance of injection over deformation) results in a decrease in the (plan view) intrusion aspect ratio; thus, high injection/low deformation results in circular intrusion, with no preferential orientation whereas low injection/high deformation results in elongated
Fig. 2. Experiment setup with contemporaneous shortening and injection for a purely brittle crust. Deformation rate is 2 cm/h and the injection rate is 20 cm/h.
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intrusions, which orient according to the main deformation features (e.g., Corti et al., 2005; Montanari et al., 2010-this volume). As an example, Fig. 3 shows two end-members behaviours observed in contractional experiments for high (80 cm/h) and low (15 cm/h) injection rates and a fixed shortening rate of 2 cm/h . These endmembers show that elongated intrusions are obtained for low injection rates (i.e., dominance of tectonic deformation over magma emplacement), whereas sub-circular intrusions (with low aspect ratio) result for high injection rates (i.e., dominance of magma emplacement over tectonics; see Montanari et al., 2010-this volume). In experiments conducted at very high injection rates, the rate of magma emplacement outpaces the rate at which the brittle host rock deforms and the shape of the intrusion is essentially governed by the expansion of magma just above the injection point. The experiment shows that the injection rate/deformation ratio of 10 used in all the performed experiments (shortening, extension and wrenching) takes into account the effects of the bulk deformation of the analogue crust on the magma emplacement. 3.2. Shortening In the shortening experiments (Fig. 4), deformation is accommodated by the thrusting and folding of the analogue upper crust. Forward propagation of deformation occurs in the presence of ductile layers, as widely described in previous experiments (e.g. Cotton and Koyi, 2000; Galland et al., 2007b). Analogue magma intrusion occurs in major contractional structures (thrust/reverse faults and at the core of anticlines), the main axis of intrusion is perpendicular to the advancing front of thrusting, roughly parallel to the hinge lines of anticlines (Fig. 4c). In both purely brittle and brittle–ductile models, the analogue magma does not reach the surface. However, in the purely brittle model, intrusion reaches higher crustal levels (∼7 cm from the base of the model, thickened due to shortening) than in the models with an embedded ductile layer (∼4 cm from the base of the model). Noteworthy, in the case of the purely brittle model, the
intruded analogue magma cut-across the main thrusts (Fig. 4b,c). In plan view, the observed spatial relationships between tectonic structures and the intruded analogue magma could thus suggest that magma intrusion occurred after the formation of thrusts (i.e. posttectonic intrusion; e.g. Castro, 1987). 3.3. Extension In the extensional experiments (Fig. 5), deformation is typically accommodated by closely-spaced normal faults defining an overall horst-and-graben-like structure. As in previous experiments (e.g. Nalpas and Brun, 1993), deformation is diffuse when a ductile layer (acting as décollement) is embedded in the sand. On the other hand, in the purely brittle model (Fig. 5) the weak analogue intrusion favoured strain localization and behaved as preferential site for normal fault propagation. In plan view, normal faults curve towards the intrusion, producing a typical hourglass pattern (e.g., Roma'n-Berdiel, 1999). In the purely brittle model, the analogue magma intrudes the entire layer of sand and reaches the surface (at a height of ∼5 cm from the base of the model). Conversely, in the brittle–ductile system, the analogue magma stops at depth (∼ 2.5 cm from the base of the model). 3.4. Strike–slip The strike–slip experiment (Fig. 6) shows a typical geometry and sequence of structure development, with early Riedel (R) shears (trending at ∼ 15° with respect to the basal Velocity Discontinuity, VD) subsequently overprinted by Y shears (sub-parallel to the basal VD) and P shears (forming at ∼ 15° with respect to the basal VD, but with an opposite trend with respect to the R shears; see Tchalenko, 1970; Naylor et al., 1986). As in previous models (e.g., Naylor et al., 1986; Corti et al., 2005), the final deformation pattern is characterized by a major zone of pervasive deformation in which anastomosing faults define shear lenses. The shape of the intrusion is strongly controlled by deformation, as the elongation of the intrusion is dictated by the
Fig. 3. Intrusion shapes resulting from shortening of a purely brittle crust with varying injection rates (15, 20, 80 cm/h) and fixed deformation rate of 2 cm/h (after Montanari et al., this volume). Upper panels show top-view photo of the intrusions; lower panels show the line drawings of the final shape of the analogue magma with details of calculation of the intrusions aspect ratio (L/W, where L and W are the intrusion width and length, respectively).
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4. Discussion
Fig. 4. Deformation and emplacement pattern in shortening experiments. (a) Top-view photo of the experiment at the end of deformation. In this case, the left-half of the model is characterized by an embedded ductile layer; the right-half is purely brittle. Note the difference in the propagation of deformation between the two halves. (b) Outlines of the major deformation geometries in the model. Grey areas indicate uplifted area above the intrusions. (c) Internal model deformation and characteristics of emplacement shown through top-view photo of the model at 2.5 cm depth (the photo has been obtained by removing the upper 2.5 cm of sand after deformation). Note the difference in the level of emplacement: in the model with embedded ductile material, intrusion occurred below this layer resulting in a deeper emplacement (∼ 4 cm from the base of the model) with respect to the purely brittle model (∼7 cm from the base of the thickened model).
sense of relative strike–slip displacement forming an angle of about 18° (brittle model) and 25° (brittle–ductile model) with respect to the basal VD (e.g., Roma’n-Berdiel et al., 1997; Corti et al., 2005). As already observed by Corti et al. (2005), the syn-deformation intrusion resulted in the uplift of the model surface in a narrow area just above the magma injection point, this fact accounts for the observed level of magma emplacement higher than the initial thickness of the model (Fig. 1). In both brittle and brittle–ductile models, the analogue magma reaches the surface. However, in the purely brittle model, surface piercing is obtained after only 6.5 cm of lateral displacement. In the model with the embedded ductile layer, the magma reaches the surface later, after about 7.5 cm of deformation.
The plot in Fig. 7 summarises the above results and clearly shows that the rheological stratification of the deforming upper crust has a major control on the depth of emplacement, especially in the case of shortening and extension. In these experiments the presence of a ductile layer in the model resulted in a deeper level of emplacement with respect to purely brittle models. During shortening, the magma did not reach the surface, even in the purely brittle model. During extension, the magma pierced the surface only in the brittle model. In strike–slip models, both purely brittle and brittle–ductile models show a similar emplacement level, as the analogue magma always reaches the surface. The level of emplacement in strike–slip models is little affected by the mechanical layering, since the main fracture zones are at high angles with and ubiquitously affect all the layers. However, the time of piercing is different in the different models: the magma rapidly reaches the surface in the brittle models, whereas piercing is delayed by the presence of a ductile layer. This implies that the presence of a soft layer controls the timing of magma emplacement. It is inferred that during strike–slip deformation of the crust, the level of magma emplacement is mainly controlled by the interplay between the bulk strain of the crust and the presence of a soft layer. As reported for magma intruded in a strike–slip setting (e.g. Corti et al., 2005) also for the contractional setting (Fig. 3) the ratio between the magma injection rate and the deformation rate mainly controls the final shape of the intrusion (see also Montanari et al., 2010-this volume). High injection rates result in magma inflation in the crust that outpaces the regional country rock deformation with nearly circular shapes of intrusion at map views (Fig. 3). In the case of shortening (Fig. 4) the absence of a weak layer strongly modifies the geometric relationships between the intruded magma and tectonic structures (thrusts and inverse faults). In fact, in the experiment with a purely brittle crust the intrusion cuts across the thrusts (discordant intrusion). The relationships of intruded magma and deformation are complex and could by simplified by two end-members (Castro, 1987; Paterson and Fowler, 1993): syn-tectonic (i.e. magma emplacement during a deformation phase) and post-tectonic (i.e. magma emplacement after the deformation phase). For the post-tectonic case, the geometric relationships between structures and intrusion in map view are quite similar to the purely brittle model in Fig. 4c. Noteworthy, very similar geometries (circular shapes and discordant intrusions) are obtained for crustal shortening when high injection rate/deformation rate ratios are used (Fig. 3). In our experiments during contraction, the magma never pierces the surface producing no volcanoes. These experimental results seem to contradict the fact that volcanoes also grew during contractional tectonic as documented by several authors (e.g. Tibaldi, 2005; Galland et al., 2007a,b; Mazzarini et al., 2008). These discrepancies could be due essentially to the limitations of our experiments to fully account for the diking processes relevant for magma transfer within the crust (e.g. Petford et al., 2000). Dykes propagation through the crust depends on the occurrence of mechanical discontinuities in the host rock. For an upwardpropagating dyke, the stiff-to-soft and the stiff-to-stiff transitions are the two most important mechanical configurations of a multilayered upper crust, as they are able to control the final level of magma emplacement (e.g. Gudmundsson, 2002; Ida, 1999). Numerical investigation (Zhang et al., 2007) considered the effect of frictional interfaces in layered rocks (bedding, parting planes). The propagation of a hydrofacture through a soft-to-stiff transition is controlled by the interface frictional strength, confining stress and fluid viscosity. High viscosity fluids (i.e. magma) induce an increase in the tensile strength of the stiff portion, so that the hydro-fracture (dyke) can penetrate the interface. However, when associated with variations in layer stiffness, the stiffto-soft transition is the most effective barrier to the upward propagation of hydro-fractures (dykes). The flexibility of the soft layer on the intact side of the interface promotes the separation and opening of
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Fig. 5. Deformation and emplacement pattern in extension experiments, illustrated as in Fig. 3 (in the schematic structural pattern, traces of normal faults are mapped at the base of the fault scarp). Again note the difference in (1) the propagation of deformation between the purely brittle and brittle/ductile models (panels a, b) and (2) the level of emplacement. In particular, in the purely brittle model the analogue magma pierced the whole sand layer and reached the surface (at an height of ∼5 cm from the base of the model); conversely, in the brittle–ductile system, the analogue magma was stopped at depth and did not reach the model surface (intrusion occurred at ∼ 2.5 cm from the base of the model). Also note the typical hourglass fault pattern in the purely brittle models due to the presence of the low-resistance intrusion that acts as a strain attractor.
the interface. This interaction is controlled by the modulus contrast between layers and by the interfacial frictional strength. The frictional interfaces are extremely important when dealing with magma moving into a deforming crust. In our experiment, although with limitations discussed above, the shear strain at the interface between stiff and soft layers (stiff-to-soft transition) promote the magma to halt its upward movement. Our experiments confirm the role of mechanical layering in controlling the level of magma emplacement and clearly show that the occurrence of soft layers in the brittle crust stops the upward movement of magma. Our experiment thus extends the control of mechanical layering to cases in which the magma has high viscosity and with magma emplacement during crustal deformation and not in pure hydrostatic conditions (e.g. Rivalta et al., 2005; Kavanagh et al., 2006). We propose that occurrence of weak (soft) layer in experiment conducted with an applied far field stress probably acts as strain attractors behaving differently than in experiments conducted in hydrostatic conditions. Future investigations will be conducted by using new material to take into account the diking processes such as vegetable oils as magma analogue (e.g. Galland et al., 2006, 2009).
Our results suggest that the occurrence of a soft layer within a brittle medium controls the level of magma emplacement in both contractional and extensional tectonic settings. These results seem to indicate that the occurrence of a soft (weak) layer controls magma emplacement, acting as an impermeable barrier to the ascent of magma (e.g. Hogan and Gilbert, 1995). Under these conditions, the rheological layering of the crust [in combination with other parameters such as the magma driving pressure and the magma injection rate] exerts a fundamental control on the level of magma emplacement and on the shape of the intrusion (Hogan et al., 1999). 5. Conclusions We presented the results of simple experiments aimed to describe the end-members conditions and their consequences on the emplacement of magma at shallow crustal level from a mechanical point of views, neglecting, as first order approximation, the thermal softening effects on host rock produced by intruding magma. The experimental results reveal that the mechanical characteristics of stratigraphic units in the upper crust strongly control the level of magma emplacement,
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Fig. 6. Deformation and emplacement pattern in strike–slip experiments, illustrated as in Figs. 2 and 3. Again note the difference in the propagation of deformation between the purely brittle and brittle–ductile models (panels a, b). The intrusion analogue makes an angle of about 18° (brittle) and 25° (brittle–ductile) with the basal VD, accordingly with the sense of shear. Also note that in both models the analogue magma pierced the sand layer and reached the surface. This occurred for low amount of lateral displacement in the purely brittle model.
in tectonic environment (i.e. compression, extension or wrenching) when small magma batches is continuously supplied. The analogue experiments show that this behaviour is straightforward in the case of crustal shortening and extension. Although the magma is emplaced at the same level in the case of strike–slip, the occurrence of a soft layer requires a higher bulk strain of the crust than in the case of a purely brittle crust. There is another important experimental finding in the case of crustal shortening: the occurrence or absence of a weak layer as well as the injection rate/deformation rate ratios strongly influences the final geometric relationships in map view among tectonic structures and coeval intrusions. The absence of a weak layer or high injection rate/deformation rate ratio led both to magma emplacement in which the final geometry of the intrusion is apparently post-tectonic. This can lead to the misinterpretation of relations between tectonics and magmatism when analysed only by
Fig. 7. Graph showing the depth of emplacement for the purely brittle and brittle/ ductile models in the different deformation settings; depth values are calculated as mean value of 3 to 4 experiments for each setting. Note that in the shortening experiment the height reached by the magma is due to the combination of magma migration and thrusting especially in the model with ductile layer.
observation at map scale. We therefore suggest that the relationship between tectonic structures and intrusive bodies in the upper crust should be considered also in light of magma intrusion rate and mechanical stratigraphy. References Albertz, M., Paterson, S.R., Okaya, D., 2005. Fast strain rates during pluton emplacement. Magmatically folded leucocratic dikes in aureoles of the Mount Stuart Batholith, Washington, and the Tuolumme Intrusive Suite, California. Geological Society of America Bulletin 117, 450–465. doi:10.1130/B25444.1. Brown, M., 1994. The generation, segregation, ascent and emplacement of granite magma. Earth Science Reviews 36, 83–130. Castro, A., 1987. On granitoid emplacement and related structures. A review. Geologische Rundschau 76, 101–124. Coleman, S.D., Gray, W., Glazner, A.F., 2004. Rethinking the emplacement and evolution of zoned plutons: geochronologic evidence for incremental assembly of the Tuolumme Intrusive Suite, California. Geology 32, 433–436. doi:10.1130/G20220.1. Cotton, J.T., Koyi, H.A., 2000. Modeling of thrust fronts above ductile and frictional detachments: application to structures in the Salt Range and Potwar Plateau, Pakistan. Geological Society of America Bulletin 112, 351–363. Corti, G., Bonini, M., Conticelli, S., Innocenti, F., Manetti, P., Sokoutis, D., 2003. Analogue modelling of continental extension: a review focused on the relations between the patterns of deformation and the presence of magma. Earth Science Reviews 63, 169–247. Corti, G., Moratti, G., Sani, F., 2005. Relations between surface faulting and granite intrusions in analogue models of strike–slip deformation. Journal of Structural Geology 27, 1547–1562. D'Lemos, R.S., Brown, M., Strachan, R.A., 1992. Granite magma generation, ascent and emplacement within a transpressional orogen. Journal of the Geological Society of London 149, 487–496. Galland, O., Cobbold, P.R., Hallot, E., de Bremond d'Ars, J., Delavaud, G., 2006. Use of vegetable oil and silica powder for scale modelling of magmatic intrusion in a deforming brittle crust. Earth and Planetary Science Letters 243, 786–804. doi:10.1016/ j.epsl.2006.01.014. Galland, O., Hallot, E., Cobbold, P.R., Ruffet, G., de Bremond d'Ars, J., 2007a. Volcanism in compressional Andean setting: a structural and geochronological study of Tromen volcano (Nequén province, Argentina). Tectonics 26, TC4010. doi:10.1029/ 2006TC002001. Galland, O., Cobbold, P.R., de Bremond d'Ars, J., Hallot, E., 2007b. Rise and emplacement of magma during horizontal shortening of the brittle crust: insights from experimental modelling. Journal of Geophysical Research 112, B06402. doi:10.1029/ 2006JB004604.
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Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o
Relations between deformation and upper crustal magma emplacement in laboratory physical models F. Mazzarini a,⁎, G. Musumeci b, D. Montanari c, G. Corti d a
Istituto Nazionale di Geofisica e Vulcanologia, Via della Faggiola 32, 56126 Pisa, Italy Dipartimento di Scienze della Terra, Università di Pisa, Via S. Maria 54, 56126, Pisa, Italy Centro di Eccellenza per la Geotermia di Larderello, Via G. Carducci, 4, 56044, Larderello (PI), Italy d Istituto di Geoscienze e Georisorse, CNR, Via G. la Pira 4, 50121 Firenze, Italy b c
a r t i c l e
i n f o
Article history: Received 19 January 2009 Received in revised form 10 June 2009 Accepted 15 September 2009 Available online 23 September 2009 Keywords: Mechanical layering of upper crust Magma emplacement Analogue modelling
a b s t r a c t This paper presents analogue models for the emplacement of granitic magmas in upper crustal levels with different mechanical layering during shortening, extension and strike–slip deformation. In particular, we investigated how a weak layer embedded in the upper brittle crust can control the level of magma emplacement. The adopted experimental setup was used to examine the control of soft rocks on the movement of magma through a deforming brittle crust. Model results indicate that the occurrence of a weak (soft) layer embedded in brittle (stiff) material has an impact on the level of magma emplacement. The level of emplacement during both extension and shortening was systematically deeper for models with a soft layer than for purely brittle models. During strike–slip deformation the magma pierced the surface in both purely brittle and brittle–ductile models. © 2009 Elsevier B.V. All rights reserved.
1. Introduction Magmatism occurs in different geodynamic settings, producing magmas that are emplaced within the crust at levels that may vary from the middle–lower crust to the surface (volcanism). In the continental crust, magma emplacement is well documented and modelled for crustal shortening (Hutton, 1997; Kalakay et al., 2001; Musumeci et al., 2005; Tibaldi, 2005; Galland et al., 2007a,b; Mazzarini et al., 2008), crustal extension (Roma'n-Berdiel, 1999; Corti et al., 2003) and wrenching (e.g. D'Lemos et al., 1992; Hutton and Reavy, 1992; Brown, 1994; Vigneresse, 1995; Roma´n-Berdiel et al., 1997; Salvini et al., 1997; Rosenberg, 2004; Corti et al., 2005). These examples document the even occurrence of intrusions (at depth in the crust) and volcanoes (at the surface) in quite different tectonic settings (convergent, divergent and strike–slip), suggesting that the crustal stress regime does not exert a first order control on the emplacement level of magma within the crust, whereas local strain distribution and near-field stress could provide room for magma emplacement. It is more likely that the level of magma emplacement is controlled by the magma supply rate and crustal rheology at least for continuous magma supply (e.g. Petford et al., 2000). This is particularly evident in the middle and upper crust, where crustal heterogeneities such as
⁎ Corresponding author. Tel.: +39 0508311956; fax +39 0508311942. E-mail address: [email protected] (F. Mazzarini). 0040-1951/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2009.09.013
lithological and rheological variations lead to mechanical weaknesses, which largely determine the depth of magma emplacement (e.g. Roma'nBerdiel et al., 1995, 1997; Kavanagh et al., 2006). The role of mechanical weaknesses seems to be more important in the upper crust, where the formation of sills often occurs by magma inflation and roof uplift, exploiting lithological variations. This is testified by the occurrence of several intrusive bodies at shallow depths within brittle–ductile rheological systems such as sedimentary covers, which contain strength anisotropies that stop the ascent of granitic magmas, allowing them to spread horizontally (Roma'n-Berdiel et al., 1995, 1997; Kavanagh et al., 2006) as well as for basaltic magmas (Pasquarè and Tibaldi, 2007; Tibaldi et al., 2008; Tibaldi and Pasquarè, 2008). The above-documented close spatial relationship between the mechanical layering of the crust and the level of magma emplacement is such that the mechanism which allows the migration of magma through a layered crust deserves further investigation. We therefore investigated the role of mechanical discontinuities in the upper brittle crust during the emplacement of magma in compressive, extensional and strike–slip regimes. To this end, analogue experiments were performed to simulate magma emplacement within a deforming brittle crust with and without a soft layer. The experimental apparatus as well as the crust and magma analogues are subsequently described. Lastly, the results of the analogue experiment are discussed, focussing on the level of emplacement and on the observed relationships between the simulated intrusions and the structures formed in the simulated deformed brittle crust.
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2. Analogue modelling We tested a series of small-scale models to investigate the influence of the mechanical layering of the upper brittle crust on the level of magma emplacement during deformation. The experiments were performed at the Tectonic Modelling Laboratory of the Institute of Geosciences and Earth Resources (National Research Council of Italy) and of the Florence University Department of Earth Sciences. Models, with initial dimensions of 60 cm × 45 cm and a thickness of 5 cm, were built above the metal base plate of the apparatus (Fig. 1). They were designed to reproduce magma intrusion both in a purely brittle crust and in an upper brittle crust comprising weak ductile layers. In the experiments, the brittle behaviour was reproduced using a dry quartzsand characterized by a grain size b250 µm, an internal friction angle of 39°, cohesion of 65 Pa and density of 1550 kg m− 3. The weak layers were reproduced using polydimethilsyloxane (PDMS), a transparent silicone with a density of 965 kg m− 3, a viscosity of 3 × 104 Pas, and Newtonian behaviour at strain rates b3 × 10− 2 s− 1 (Weijermars, 1986). A low viscosity mixture of silicone (Wacker Silicone Bouncing Putty 29) and oleic acid (4:1% in weight) was employed to simulate granitic magma; this mixture exhibits Newtonian behaviour (with viscosity of 7 × 102 Pa s) and a density of 1060 kg m− 3. We performed three different experimental series representing three different tectonic regimes (strike–slip, shortening and extension); for each tectonic regime we tested two models constructed with and without the embedded ductile layer. In the case of shortening, we also tested a
model in which one half has a purely brittle rheology and the other has an embedded ductile layer. Each half has an injection point located at the centre of the base of the model. A moving wall driven by a stepper motor controlled by a central unit enabled deformation of the models (Fig. 1a). Magma intrusion during deformation was simulated using a special injection apparatus consisting of a piston (Fig. 1b) and a magma distribution system made of pipes with a central injection point (10 mm in diameter) on the metal base plate of the apparatus (Fig. 1c). The deformation rate (2 cm/ h) and the magma injection velocity (20 cm/h) remained constant in all the experiments. 2.1. Scaling The experiments were scaled taking into consideration the conditions of geometric, dynamic, kinematic and rheologic similarity outlined by Hubbert (1937), Ramberg (1981) and Weijermars and Schmeling (1986). Geometric similarity was achieved by imposing a scaling ratio of 10− 5, such that 1 cm in the experiments corresponds to 1 km in nature. Density of modelling materials imposes a density scale ratio of ∼0.4–0.5; together with the reduction factor of density and gravity (=1), the resulting model to nature ratio of stresses was of the order of 4 × 10− 6. The scaling ratio of time is ∼1.5 × 10− 9 (1 h in the models represents ∼ 80,000 yrs in nature) and the resulting scaling factor of the linear injection velocity is ∼6.6 × 103. The scaling ratio of viscosity (∼ 6 × 10− 15) results in a scaled viscosity of ∼5 × 1018 Pa s for
Fig. 1. Details of the experimental set up. (a) Oblique view of the deformation apparatus; (b) analogue magma injection piston; (c) metal base of the apparatus showing the magma injection point and the moving wall/fixed base (note that the moving plate attached to the moving wall is used in strike–slip experiments only); (d) schematic model cross-section showing the rheological stratification of models. In the experiments, the modelling length ratio is l⁎ = 1 × 10− 5 (1 cm in the model represents about 1 km in nature).
F. Mazzarini et al. / Tectonophysics 484 (2010) 139–146 Table 1 Main scaling parameters (see Corti et al., 2005). Parameter
Model
Nature
Model / Nature
Length (m) Viscosity (ductile layer) (Pa s) Viscosity (magma) (Pa s) Displacement velocity Emplacement velocity
0.01 3000 700 2 (cm h− 1) 20 (cm h− 1)
1000 ∼5 × 1018 Pa s ∼1017 Pa s 2.7 (cm yr− 1) ∼30 (cm yr− 1)
1 × 10− 5 ∼6 × 10− 15 ∼6 × 10− 15 ∼6.6 × 103 ∼6.6 × 103
the weak ductile layer and ∼ 1017 Pa s for the granitic magma. Scaling parameters for viscosity and for the magma input rate and deformation are reported in Table 1. Details of the scaling procedure are reported in Montanari et al. (2010-this volume). 2.2. Limitations of modelling The performed analogue experiments are only an approximation of nature, and the following should be borne in mind: (i) only simple, purely convergent, divergent or strike–slip tectonic settings were investigated; (ii) the scaling is based on mechanical principles, not thermal ones. In our models, we assume isothermal conditions and do not take into account the thermal effects of intrusion-related heating and pluton cooling on the rheology of the crust and magma, respectively. This assumption is based on the fact that the magma-related thermal perturbation (contact aureole) affects small portions of the host rocks and that magmas may not freeze quickly in actively deforming regions, owing to persisting thermal anomalies related to the repeated intrusion of magmas into the crust. We are aware that, as consequence of this assumption, in our experiment the deformation in the analogue crust is essentially accomplished by faulting and bending without formation of cleavage and/or ductile deformation because we are not able to reproduce the formation of a thermal aureole that could accommodate the high strain induced in the host rock by the magma emplacement (e.g. Albertz et al., 2005); (iii) although the low viscosity mixture of silicone and oleic acid reproduces a significant viscosity contrast between the low viscosity melt fraction and the crust, when scaled to nature, it is only able to simulate a high viscosity magma. This implies that the analogue magma does not move upward through hydrofracturing alone, but also in diapirs. A section of an experiment performed in compression on a purely brittle modelled crust is shown in Fig. 2. The analogue magma rises along the main thrust and locates in the core of the push-up structure where local strain creates room for magma. The adopted material and scaling thus determine magma movement in our experiment by pervasive flow through shear zones and diapirims. Nonetheless, we consider that in these experiments the control of
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mechanical layering on the magma emplacement is always accounted for; (iv) as each layer in the experimental setup is homogeneous, the role of inherited tectonic structures is neglected. Notwithstanding the high mechanical heterogeneity of the upper crust, this assumption is acceptable because we mainly aimed to investigate the level of magma emplacement as a function of interplay between mechanical layering and tectonic regime. (v) magma emplacement in a deforming crust is the result of the interaction of several processes (magma emplacement and cooling, metamorphism and deformation) having different time scales and rates (e.g. Paterson and Tobish, 1992). The injection rates (i.e. the rate at which the magma is introduced in the system), once properly scaled, are in the 20–106 cm yr− 1 range. The scaled deformation rate used in the experiments, is 2.7 cm yr− 1, and the resulting scaled magma supply is in the range 10− 2–10− 4 km3 yr− 1. These values are compatible with rates of faulting, magma supply and magma emplacement observed in nature (Paterson and Tobish, 1992 and references therein; Jellinek and DePaolo, 2003). (vi) the rate at which pluton grows is a critical point. Large batholiths, for example the Tuolumme Intrusive Suite in California, grew by incremental assemblage of small magma batches over millions years (Coleman et al., 2004; Glazner et al., 2004). On the other hands it has been documented that ascent of magma trough the crust and magma emplacement could be a very fast processes (e.g. Petford et al, 2000). In our experiment we consider continuous magma supply of relatively small magma batches (low magma supply, see point above) at very shallow crustal level. 3. Experimental results For each experiment, deformation was monitored through direct observation and sequential top-view photos. At the end of each experiment, the model was sectioned and the level of magma emplacement (height of the magma above the base of the model) was measured. As magma reaches the model surface the height of the surface has been used as the level of magma emplacement. In all the experiments the lowermost boundary of analogue intrusions is at the base of the model, that is the level where the injection point is located. 3.1. Intrusion shape For each deformation setting, it is well known that the shape of the intrusion is primarily controlled by the ratio between the intrusion and deformation velocity (e.g. Roma'n-Berdiel et al., 1997; Corti et al., 2005; Montanari et al., 2010-this volume). An increase in this ratio (i.e., predominance of injection over deformation) results in a decrease in the (plan view) intrusion aspect ratio; thus, high injection/low deformation results in circular intrusion, with no preferential orientation whereas low injection/high deformation results in elongated
Fig. 2. Experiment setup with contemporaneous shortening and injection for a purely brittle crust. Deformation rate is 2 cm/h and the injection rate is 20 cm/h.
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intrusions, which orient according to the main deformation features (e.g., Corti et al., 2005; Montanari et al., 2010-this volume). As an example, Fig. 3 shows two end-members behaviours observed in contractional experiments for high (80 cm/h) and low (15 cm/h) injection rates and a fixed shortening rate of 2 cm/h . These endmembers show that elongated intrusions are obtained for low injection rates (i.e., dominance of tectonic deformation over magma emplacement), whereas sub-circular intrusions (with low aspect ratio) result for high injection rates (i.e., dominance of magma emplacement over tectonics; see Montanari et al., 2010-this volume). In experiments conducted at very high injection rates, the rate of magma emplacement outpaces the rate at which the brittle host rock deforms and the shape of the intrusion is essentially governed by the expansion of magma just above the injection point. The experiment shows that the injection rate/deformation ratio of 10 used in all the performed experiments (shortening, extension and wrenching) takes into account the effects of the bulk deformation of the analogue crust on the magma emplacement. 3.2. Shortening In the shortening experiments (Fig. 4), deformation is accommodated by the thrusting and folding of the analogue upper crust. Forward propagation of deformation occurs in the presence of ductile layers, as widely described in previous experiments (e.g. Cotton and Koyi, 2000; Galland et al., 2007b). Analogue magma intrusion occurs in major contractional structures (thrust/reverse faults and at the core of anticlines), the main axis of intrusion is perpendicular to the advancing front of thrusting, roughly parallel to the hinge lines of anticlines (Fig. 4c). In both purely brittle and brittle–ductile models, the analogue magma does not reach the surface. However, in the purely brittle model, intrusion reaches higher crustal levels (∼7 cm from the base of the model, thickened due to shortening) than in the models with an embedded ductile layer (∼4 cm from the base of the model). Noteworthy, in the case of the purely brittle model, the
intruded analogue magma cut-across the main thrusts (Fig. 4b,c). In plan view, the observed spatial relationships between tectonic structures and the intruded analogue magma could thus suggest that magma intrusion occurred after the formation of thrusts (i.e. posttectonic intrusion; e.g. Castro, 1987). 3.3. Extension In the extensional experiments (Fig. 5), deformation is typically accommodated by closely-spaced normal faults defining an overall horst-and-graben-like structure. As in previous experiments (e.g. Nalpas and Brun, 1993), deformation is diffuse when a ductile layer (acting as décollement) is embedded in the sand. On the other hand, in the purely brittle model (Fig. 5) the weak analogue intrusion favoured strain localization and behaved as preferential site for normal fault propagation. In plan view, normal faults curve towards the intrusion, producing a typical hourglass pattern (e.g., Roma'n-Berdiel, 1999). In the purely brittle model, the analogue magma intrudes the entire layer of sand and reaches the surface (at a height of ∼5 cm from the base of the model). Conversely, in the brittle–ductile system, the analogue magma stops at depth (∼ 2.5 cm from the base of the model). 3.4. Strike–slip The strike–slip experiment (Fig. 6) shows a typical geometry and sequence of structure development, with early Riedel (R) shears (trending at ∼ 15° with respect to the basal Velocity Discontinuity, VD) subsequently overprinted by Y shears (sub-parallel to the basal VD) and P shears (forming at ∼ 15° with respect to the basal VD, but with an opposite trend with respect to the R shears; see Tchalenko, 1970; Naylor et al., 1986). As in previous models (e.g., Naylor et al., 1986; Corti et al., 2005), the final deformation pattern is characterized by a major zone of pervasive deformation in which anastomosing faults define shear lenses. The shape of the intrusion is strongly controlled by deformation, as the elongation of the intrusion is dictated by the
Fig. 3. Intrusion shapes resulting from shortening of a purely brittle crust with varying injection rates (15, 20, 80 cm/h) and fixed deformation rate of 2 cm/h (after Montanari et al., this volume). Upper panels show top-view photo of the intrusions; lower panels show the line drawings of the final shape of the analogue magma with details of calculation of the intrusions aspect ratio (L/W, where L and W are the intrusion width and length, respectively).
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4. Discussion
Fig. 4. Deformation and emplacement pattern in shortening experiments. (a) Top-view photo of the experiment at the end of deformation. In this case, the left-half of the model is characterized by an embedded ductile layer; the right-half is purely brittle. Note the difference in the propagation of deformation between the two halves. (b) Outlines of the major deformation geometries in the model. Grey areas indicate uplifted area above the intrusions. (c) Internal model deformation and characteristics of emplacement shown through top-view photo of the model at 2.5 cm depth (the photo has been obtained by removing the upper 2.5 cm of sand after deformation). Note the difference in the level of emplacement: in the model with embedded ductile material, intrusion occurred below this layer resulting in a deeper emplacement (∼ 4 cm from the base of the model) with respect to the purely brittle model (∼7 cm from the base of the thickened model).
sense of relative strike–slip displacement forming an angle of about 18° (brittle model) and 25° (brittle–ductile model) with respect to the basal VD (e.g., Roma’n-Berdiel et al., 1997; Corti et al., 2005). As already observed by Corti et al. (2005), the syn-deformation intrusion resulted in the uplift of the model surface in a narrow area just above the magma injection point, this fact accounts for the observed level of magma emplacement higher than the initial thickness of the model (Fig. 1). In both brittle and brittle–ductile models, the analogue magma reaches the surface. However, in the purely brittle model, surface piercing is obtained after only 6.5 cm of lateral displacement. In the model with the embedded ductile layer, the magma reaches the surface later, after about 7.5 cm of deformation.
The plot in Fig. 7 summarises the above results and clearly shows that the rheological stratification of the deforming upper crust has a major control on the depth of emplacement, especially in the case of shortening and extension. In these experiments the presence of a ductile layer in the model resulted in a deeper level of emplacement with respect to purely brittle models. During shortening, the magma did not reach the surface, even in the purely brittle model. During extension, the magma pierced the surface only in the brittle model. In strike–slip models, both purely brittle and brittle–ductile models show a similar emplacement level, as the analogue magma always reaches the surface. The level of emplacement in strike–slip models is little affected by the mechanical layering, since the main fracture zones are at high angles with and ubiquitously affect all the layers. However, the time of piercing is different in the different models: the magma rapidly reaches the surface in the brittle models, whereas piercing is delayed by the presence of a ductile layer. This implies that the presence of a soft layer controls the timing of magma emplacement. It is inferred that during strike–slip deformation of the crust, the level of magma emplacement is mainly controlled by the interplay between the bulk strain of the crust and the presence of a soft layer. As reported for magma intruded in a strike–slip setting (e.g. Corti et al., 2005) also for the contractional setting (Fig. 3) the ratio between the magma injection rate and the deformation rate mainly controls the final shape of the intrusion (see also Montanari et al., 2010-this volume). High injection rates result in magma inflation in the crust that outpaces the regional country rock deformation with nearly circular shapes of intrusion at map views (Fig. 3). In the case of shortening (Fig. 4) the absence of a weak layer strongly modifies the geometric relationships between the intruded magma and tectonic structures (thrusts and inverse faults). In fact, in the experiment with a purely brittle crust the intrusion cuts across the thrusts (discordant intrusion). The relationships of intruded magma and deformation are complex and could by simplified by two end-members (Castro, 1987; Paterson and Fowler, 1993): syn-tectonic (i.e. magma emplacement during a deformation phase) and post-tectonic (i.e. magma emplacement after the deformation phase). For the post-tectonic case, the geometric relationships between structures and intrusion in map view are quite similar to the purely brittle model in Fig. 4c. Noteworthy, very similar geometries (circular shapes and discordant intrusions) are obtained for crustal shortening when high injection rate/deformation rate ratios are used (Fig. 3). In our experiments during contraction, the magma never pierces the surface producing no volcanoes. These experimental results seem to contradict the fact that volcanoes also grew during contractional tectonic as documented by several authors (e.g. Tibaldi, 2005; Galland et al., 2007a,b; Mazzarini et al., 2008). These discrepancies could be due essentially to the limitations of our experiments to fully account for the diking processes relevant for magma transfer within the crust (e.g. Petford et al., 2000). Dykes propagation through the crust depends on the occurrence of mechanical discontinuities in the host rock. For an upwardpropagating dyke, the stiff-to-soft and the stiff-to-stiff transitions are the two most important mechanical configurations of a multilayered upper crust, as they are able to control the final level of magma emplacement (e.g. Gudmundsson, 2002; Ida, 1999). Numerical investigation (Zhang et al., 2007) considered the effect of frictional interfaces in layered rocks (bedding, parting planes). The propagation of a hydrofacture through a soft-to-stiff transition is controlled by the interface frictional strength, confining stress and fluid viscosity. High viscosity fluids (i.e. magma) induce an increase in the tensile strength of the stiff portion, so that the hydro-fracture (dyke) can penetrate the interface. However, when associated with variations in layer stiffness, the stiffto-soft transition is the most effective barrier to the upward propagation of hydro-fractures (dykes). The flexibility of the soft layer on the intact side of the interface promotes the separation and opening of
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Fig. 5. Deformation and emplacement pattern in extension experiments, illustrated as in Fig. 3 (in the schematic structural pattern, traces of normal faults are mapped at the base of the fault scarp). Again note the difference in (1) the propagation of deformation between the purely brittle and brittle/ductile models (panels a, b) and (2) the level of emplacement. In particular, in the purely brittle model the analogue magma pierced the whole sand layer and reached the surface (at an height of ∼5 cm from the base of the model); conversely, in the brittle–ductile system, the analogue magma was stopped at depth and did not reach the model surface (intrusion occurred at ∼ 2.5 cm from the base of the model). Also note the typical hourglass fault pattern in the purely brittle models due to the presence of the low-resistance intrusion that acts as a strain attractor.
the interface. This interaction is controlled by the modulus contrast between layers and by the interfacial frictional strength. The frictional interfaces are extremely important when dealing with magma moving into a deforming crust. In our experiment, although with limitations discussed above, the shear strain at the interface between stiff and soft layers (stiff-to-soft transition) promote the magma to halt its upward movement. Our experiments confirm the role of mechanical layering in controlling the level of magma emplacement and clearly show that the occurrence of soft layers in the brittle crust stops the upward movement of magma. Our experiment thus extends the control of mechanical layering to cases in which the magma has high viscosity and with magma emplacement during crustal deformation and not in pure hydrostatic conditions (e.g. Rivalta et al., 2005; Kavanagh et al., 2006). We propose that occurrence of weak (soft) layer in experiment conducted with an applied far field stress probably acts as strain attractors behaving differently than in experiments conducted in hydrostatic conditions. Future investigations will be conducted by using new material to take into account the diking processes such as vegetable oils as magma analogue (e.g. Galland et al., 2006, 2009).
Our results suggest that the occurrence of a soft layer within a brittle medium controls the level of magma emplacement in both contractional and extensional tectonic settings. These results seem to indicate that the occurrence of a soft (weak) layer controls magma emplacement, acting as an impermeable barrier to the ascent of magma (e.g. Hogan and Gilbert, 1995). Under these conditions, the rheological layering of the crust [in combination with other parameters such as the magma driving pressure and the magma injection rate] exerts a fundamental control on the level of magma emplacement and on the shape of the intrusion (Hogan et al., 1999). 5. Conclusions We presented the results of simple experiments aimed to describe the end-members conditions and their consequences on the emplacement of magma at shallow crustal level from a mechanical point of views, neglecting, as first order approximation, the thermal softening effects on host rock produced by intruding magma. The experimental results reveal that the mechanical characteristics of stratigraphic units in the upper crust strongly control the level of magma emplacement,
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Fig. 6. Deformation and emplacement pattern in strike–slip experiments, illustrated as in Figs. 2 and 3. Again note the difference in the propagation of deformation between the purely brittle and brittle–ductile models (panels a, b). The intrusion analogue makes an angle of about 18° (brittle) and 25° (brittle–ductile) with the basal VD, accordingly with the sense of shear. Also note that in both models the analogue magma pierced the sand layer and reached the surface. This occurred for low amount of lateral displacement in the purely brittle model.
in tectonic environment (i.e. compression, extension or wrenching) when small magma batches is continuously supplied. The analogue experiments show that this behaviour is straightforward in the case of crustal shortening and extension. Although the magma is emplaced at the same level in the case of strike–slip, the occurrence of a soft layer requires a higher bulk strain of the crust than in the case of a purely brittle crust. There is another important experimental finding in the case of crustal shortening: the occurrence or absence of a weak layer as well as the injection rate/deformation rate ratios strongly influences the final geometric relationships in map view among tectonic structures and coeval intrusions. The absence of a weak layer or high injection rate/deformation rate ratio led both to magma emplacement in which the final geometry of the intrusion is apparently post-tectonic. This can lead to the misinterpretation of relations between tectonics and magmatism when analysed only by
Fig. 7. Graph showing the depth of emplacement for the purely brittle and brittle/ ductile models in the different deformation settings; depth values are calculated as mean value of 3 to 4 experiments for each setting. Note that in the shortening experiment the height reached by the magma is due to the combination of magma migration and thrusting especially in the model with ductile layer.
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