Crustal regime of deformation and ascent of granitic magma

TECTONOPHYSICS Tectonophysics249 (1995) 187-202

ELSEVIER

Crustal regime of deformation and ascent of granitic magma Jean Louis Vigneresse

1

CREGU, Vandoeuvre. France

Received 14 February 1994; accepted 11 January 1995

Abstract Structural studies and gravity data demonstrate the magma flow pattern of granitic plutons and their morphology at depth. The deepest zones calculated from gravity data, provided they show vertical lineations in outcrop, are interpreted as the feeder channels which fed the pluton. Their morphology at depth, as well as the position and number of root zones, vary according to the type of deformation, suggesting that regional deformation determines the emplacement of granitic plutons. Therefore, deformation must also influence the segregation and the ascent of acid magmas. A thermal model is constructed which demonstates the structure and heat generation of the Hercynian crust. According to that model, the thermal gradient changes the conditions for the deformation of the crust. Obviously, a large amount of heat must be supplied to induce melting of the lower crust. As a result magma segregation and ascent occur entirely within the ductile field. When upwelling magma reaches the brittle/ductile transition zone, the hydrodynamic force (buoyancy) is insufficient to fracture the upper brittle crust. Parameters which control diapirism (wavelength, viscosity, thickness) do not match those related to the volume of erupted magma. As a consequence, diapiric upwelling appears to be limited under favourable conditions to the lower ductile crust and does not apply to intrusions in a brittle crust. Shearing helps to segregate magma and to feed conduits toward the surface since it creates the required zones of local extension for the magma to infiltrate.

1. I n t r o d u c t i o n Granitic plutons result from melting, segregation, upwelling and crystallization of acid magmas at higher levels in the continental crust. Such intrusions are studied by direct field observations or by laboratory modeling. Analogue and numerical models have proven that R a y l e i g h - T a y l o r instabilities develop with large viscosity and density contrasts, resulting in diapiric upwelling of the more mobile material (Ramberg, 198l). Field observations fail to show acid magma movement. Therefore only the final

J Also at ENS Grologie, Nancy, France.

stages of magma emplacement can be observed. Extrapolation is required to explain mechanisms which may have occurred at deeper levels and during the time the magma was emplaced. Popular ideas such as diapirs can induce biased concepts in textbook readers' minds due to the fact that cartoons are not to scale. As an example, I show the "schematic representation of an igneous diapir", redrawn from a textbook (Price and Cosgrove, 1990). The cartoon is shown at the scale of the textbook as true to scale (Fig. 1). To adjust the scale, I suggest that contact metamorphism occurs at about 300°C (approx. 7 km) and that melting occurs at a depth of approx. 20 km, or 600 MPa (6 kbar). The pluton is 30 km in radius. The correct scale of the cartoon

0040-1951/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSD! 0040- 195 1(95)00005-4

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J.L. Vigneresse / Tectomq~hysics 249 (1995) 187-202

completely changes its dimensions. The popular inverted tear drop with vertical long axis, underlying all diapiric concepts, turns into a flattened cupola. I propose to scrutinize granite intrusions from this novel point of view. Few data are available in order to quantify the ascent mechanism of acid magmas. Most currently available data relate to the structure and the shape of the granitic bodies observed in the field or indirectly at depth by geophysical methods. They stem from many surveys (Guillet et al., 1985; Guineberteau et al., 1987; Audrain et al., 1989a, b; B r u n e t al., 1990) mostly carried out over biotite, or biotite-cordierite granites, the most evolved being the two-mica granites. Consequently, the present model is limited to those types of magma and does not apply to ring-dyke complexes, nor to mantle-derived magmas, which may behave differently. Other parameters which control granite emplacement are the thermal and deformation conditions in the crust. The latter depends on the nature and thermal state of the crust. Present crustal structure

surface

and petrological composition determine the range of available heat generation, and control the thermal conditions in the crust. The pressure and temperature conditions assigned to the source area of melting granite must fit the data proposed by experimental petrologists. To produce granitic melts, a large amount of heat must be added to a thermally stable crust. When a convenient thermal state yielding a granitic melt is established for the crust, the theologic behaviour of the latter can be calculated. Deformation controls the final stage of granite emplacement in the crust (Vigneresse, 1995). It is therefore certainly active during the collection and ascent of magma. During this stage, no new data are presented on this part of the ascent. A brief review of existing possibilities is presented, with emphasis given to the role played by delormation. However, the morphology and volume of plutons place constraints on the segregation and ascent of acid magmas in a theologically defined crust. In particular, 1 propose that diapiric processes should not be considered as an efficient mechanism to enable granitic

30 km

30 km

enclaves 7

\

%-/ -

20

-

30 km

]

J

]

partial

/

~mng

ca 20kin Fig. I. Cartoon, redrawn after Price and Cosgrove (1990) showing, on the left, a "schematic representation of an igneous diapir", but unfortunatelynot to scale, with, on the right, the same structure using average values for its dimensions.Contact metamorphismis assumed to take place at 300°C, or around 7 kin, and partial melting occurs at approx. 20 km, thus correspondingto the triple points for silicates, around 600 MPa. The radius of the pluton is approx. 30 kin. I strongly propose that all cartoonistsgive some thought to the scale they use because of the fallacious concepts the picture may convey.

J.L. Vigneresse / Tecmnophysics 249 (1995) 187-202

magmas to intrude the upper continental crust. Data on the spacing of plutons, volume of intruded magmas, as well as geophysical information on the shape at depth of plutons do not favour a diapiric ascent. Channeling of magma through small conduits is thought more likely.

2. Structure and root z o n e s in granites

The magma flow pattern in a still viscous magma is recorded by the first crystallized minerals. They define foliation planes and lineations. Structural studies on granites (Guillet et al., 1985; Guineberteau et al., 1987) or techniques using the anisotropy of the magnetic susceptibility (Bouchez et al., 1990) assist in mapping this pattern. To determine deeper structures, geophysical data and among them gravity surveys enable the estimation of the shape ofthe plutons (Vigneresse, 1990). Measurements must be taken on a grid similar to geological observations in order to correlate both types of information. Geological variations (facies changes) must be taken into account for possible changes in density (Vigneresse, 1990). Depending on the density coverage and on the accuracy of the measurements, the precision of depth values is about 15-20%, because of the confidence loss in estimating the density at depth. The shape of the floor is determined with better precision, about 5%, since it reflects the homogeneity and the density of the sample stations. A combination of both structural and geophysical data is required to fully understand the mode of emplacement of a pluton (Guillet et al., 1985; Guineberteau et al., 1987; Audrain et al., 1989a, b; Brun et al., 1990). When the erosion level of the pluton only affects its uppermost part, the magmatic structures reflect the interaction between the emplacement of the pluton and the surrounding rocks (Hanmer and Vigneresse, 1980; Paterson et al., 1989). For a deeper eroded pluton, zones, where magmatic lineations are vertical, often correlate with the deepest zones calculated from gravity data, where the dip of the floor strongly increases. Conversely, zones where lineations are weakly dipping often denote shallow to medium depth. The deeper zones with vertical lineations are interpreted as the magma feeder channels. They are apparently active during the en-

189

tire time the pluton has been emplaced and are often characterized at the outcrop by more evolved magma or late intrusive facies. They are located between 5 and 14 km at depth and have a moderate surface area of 1 to 10 km 2, compared to several hundreds of square kilometers for the surface area of a pluton. The geometry and orientation of the root zones yield direct information on the rheological properties of the crust. They also reflect how the magma has been intruded controlled by the regional stress field (Vigneresse, 1995). The depth at which the root zones are presently observed must be corrected for the effect of erosion. This is achieved using the temperature of the contact metamorphism around the intrusion.

3. Root zones in granites and the field of d e f o r m a tion

The geometry of granitic bodies and its relationship to deformation have been presented in a companion paper (Vigneresse, 1995). Here I only summarize the major results. Granitic plutons intrusive during shear stress conditions (Fig. 2) are not very thick (about 5 _+ 2 km) and contain only one feeder, two in larger plutons, extending down to a maximum depth of 5 - 1 4 kin. In contrast, granites intrusive during extensional stress conditions show a very thin (about 2 - 3 km) main unit with subhorizontal magmatic foliation planes and horizontal lineations intruded by variously evolved facies in which magmatic structures are subvertical (Fig. 2). Granites emplaced during regional compression generally have a moderate volume. Plutons emplaced in between overlapping faults result from the infilling of magma as soon as deformation creates open spaces. Consequently, they have steep walls and a flat floor. Whatever the type of deformation, the roots occur within the local extensional area relative to the regional stress field. Geometrically, two situations are observed. One, the less common, corresponds to a root aligned with the major stress component (cr~), reflecting intrusion under brittle conditions. In the other, the root intersects o% at a high angle and is aligned with the extensional stress. This reflects plastic behavior during which stresses reorientate the plastic body towards the extensional direction. In

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

190

--1

Extension

1

A

A,

i

13

3'

_c

~~

.

Overlapping shear

G~ ®

~

G,

cFig. 2. Schematic morphology of granitic plutons according to the regional field of deformation, with dotted areas showing the root zones. A

map (on the left) and a cross section (on the right) with no vertical exaggeration are presented for: A, extension; B, wrench deformation; C, opening fractures during shear deformation. The figures are highly simplified from real plutons taken as examples (Vigneresse, 1995).

both cases, the active role played by deformation during granite emplacement (Pitcher, 1979) is being emphasized.

4. Volume and geometry of plutons Gravity data have been collected from many granitic plutons, emplaced during varied tectonic conditions (Vigneresse, 1988a, 1995, and references). Most relevant information describing the morphology of plutons has been incorporated into a data bank in order to delineate the major trends. Because all surveys have been interpreted in three dimensions, the distribution of the m a g m a volume with depth is known (Fig. 3). The relationship between the radius of a cone and the volume of magma has also been calculated. This shows the variation of cone aperture with depth (Fig. 4).

One major result from the present data relates to the total volume of intruded xnagma. On average about 500 to 1000 km 3 of magma erupted through a single root. Some plutons are larger, but contain several roots, two for Cabeza de Araya and Pontivy and many for Saint Sylvestre or La Marche. Some other large plutons, like those in the Fichtelgebirgc or the Vend6e, are in fact composite plutons which makes it difficult to separate the respective volume of each batch of magma. Because presently not all plutons have the same erosion level, the volumes are certainly shifted towards minimum values. Therefore, the maximum volume of magma delivered by one root could range between 1000 and 1500 km 3, corresponding to a pluton of about 400 km 2 in outcrop. The cone angle varies with depth (Fig. 4). In the uppermost crust, the floor plunges at 30 °, and gradually steepens to 45 ° at a moderate depth ( 3 - 6 km). Lower in the crust, the cone sharpens and its

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202 VOLUME 0

50o

1000

m

m

dip reaches about 60 ° . The change at depth between a gradually increasing plunge to a steeper one reflects the transition zone between areas of brittle and ductile deformation. With this data set, the model can easily be transferred from one pluton to another (Fig. 4) taking erosion into account. In particular, no difference is observed in the volume distribution with depth between plutons emplaced during extension and plutons emplaced during shear or compression, in spite of their pronounced differences in morphology. An explanation may be the mode of magma distribution at each level. For plutons emplaced during shear or compression, one root delivers the magma, the volume of which strongly decreases with depth. Conversely, in plutons emplaced during extension, the larger number of roots delivers a larger amount of magma.

1500

SAINT-SYL... L A MARCHE LOCRONAN PONTIVY

____ Z I--

GUEHENNO L A GACILLY

____

E3

SAINT-RE... PLOUARET FLAMANVI...

10

HUELGOAT MORTAGNE __

12

_ VENDEE

____

191

FICHTELGE... C A S E Z A D... GUITIRIZ

14

HOMBREIRO

5. T h e r m a l model of the crust Fig. 3. Volumeof granitic plutons (in km3) at each depth level (in kin) as interpreted from gravity data. The following data encompass: Saint-Sylvestreand La Marche, both in the French Massif Central; Locronan, Pontivy, Gu6henno, La Gacilly in southern Brittany; Saint-Renan, Plouaret, Huelgoat, in northern Brittany; Mortagne, Vend6e in southeastern Brittany, France; Flamanville in Normandy,France; Fichtelgebirgein Bayern, Germany;Cabeza de Araya in Estremadura, Spain; Guitiriz, Hombreiro in Galicia, Spain. All plutons are Hercynian. EQUIVALENT

0

5

21

A thermal model has been constructed to examine the physical conditions leading to segregation and collection of acid magma before it wells up (Fig. 5). The model presented is static, in contrast to a dynamic one in which the ascent of magma is essential. The thermal gradient must reach the temperature required to initiate melting at the depth of the source RADIUS

I0

15

20

25

30

6

J~//i O 10

~

LA M A R C H E

LOCRONAN ___ PONTIVY

_~//¢

4 1 4 I I# 14

¢

41"// ~ //

FLAMANVILLE ---

MORTAGNE VENDEE

-------GUEHENNO

----

FICHTELGEBIRGE

- -

- -

CABEZA DE AR...

GAC,LLY

--SAINT-RENAN -- -- -- PLOUARET

GUITIRIZ HOMBREIRO

Z

Fig. 4. Values(in km) of the radius of a cone with similar volume distributionat depth as in granitesfrom Fig. 3. Same data set with similar symbols are used. Horizontalscale is equal to the vertical one.

192

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

area. These conditions can be deduced from experiments on partial melting of pelitic or amphibolitic material (Vielzeuf and Holloway, 1988; Rushmer, 1991; Patino-Douce and Johnston, 1991). They vary according to the amount and composition of the available fluid. Starting which wet pelites, the conditions which produce two-mica granites are estimated at 700°C and 600 MPa, whereas amphibolites melt at 850-900°C under dry conditions. From this model, the conditions of deformation prevailing during the various stages o f the ascent can be deduced. The structure of the Hercynian crust is adopted as the standard for a stable continental crust since its thermal properties and composition are well determined (Lucazeau and Vasseur, 1981; Jolivet et al., 1989). It is represented as a three-layered model from seismic measurements (BIRPS and ECORS, 1986). A lower layer with a seismic velocity of 6.8 k m / s is overlain by a 10-kin-thick layer with a velocity o f 6.2 k i n / s , and then by a 8-km-thick layer with a velocity of 5.8 k m / s . The lowermost layer can be assigned to the granulitic facies based on its seismic velocity. The intermediate layer is inter-

preted as being amphibolitic, whereas the uppermost layer is granodioritic. This is a highly simplified model of a more complex crust, but it explains the composition of the crust better than any other model of heat generation and flow with depth. Crustal xenoliths (Lucazeau and Vasseur, 1981; Vigneresse, 1988b) indicate a heat generation of 0.2 / z W / m ~ in the granulitic facies, and values of approx. 1.7 p , W / m 3 for amphibolites. The average heat generation of the upper crustal rocks is approx. 2 . 9 / x W / m 3 (Jolivet et al., 1989). Hercynian granites have been omitted from this model (Fig. 5) for simplicity, but their effect is not relevant to the thermal state of the lower crust where granites were formed. The pressure in the source area (above 600 MPa) corresponds to a depth of approx. 17 km assuming lithostatic pressure. According to the proposed model, it also corresponds to the base of the amphibolitic layer, which certainly melts before the granulites. In a stable crust with an average gradient of 2 0 ° C / k i n , the temperature at this depth is 350°C. Therefore, a large quantity of heat must be added to produce granitic melts. The thermal gradient steepens either

1 O0 200 M P a

T °C

500

granodiorite 2.9 gW/m 3

10 . . . .

20

30

~i Km

~ DB

i+\~ i"mx~~ ~ 1 S

B D

amphibolite

B

1.7 W/m granulite

0.2 ~W/m3

t4t

S Y (

i/

Km

Fig. 5. Hercyt-lan crustal model constructed from geochemical, geophysical and structural data. On the left, the geothermal gradient is calculated according to the heat generation (A) distribution with depth and heat flow (Q) coming from the mantle. Two situations are presented. One (in gray), with a low basal heat flow of 15 mW/m 2 corresponds to a stable crust. The second (in lighter gray) with a large heat flow (65 mW/m 2) is that of a thermally perturbated crust in which melt develops. It serves to determine, on the right, the differential strength under extensional conditions. The brittle/transition zone is dotted.

193

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

0

,

200

400

,

600

800

M

Pa

or1- (r3 \E

~

38°/kin

10

20

Fig. 6. Strain versus depth calculated during extensional (E), shearing (S) and compressional (C) stress conditions. The brittle/ductile transition is displayed according to three different values of the geothermal gradient. The hydrodynamicforces of a magma column starting at the Moho is drawn on the left to show the differential strain required to break the brittle crust.

because heat is added from below, or because the crust is thinner. In the present model, there is no specific reason to involve large-scale crustal thinning, and heat is simply added from below. A temperature higher than 700°C at 17 km implies a thermal gradient of 4 0 ° C / k m . In a conductive model, this is achieved by increasing the basal heat flow to 65 m W / m 2. Such a value is about four times the heat flow coming from the continental mantle in stable areas. It yields a surface thermal gradient of 4 0 ° C / k m and a surface heat flow of 130 m W / m 2 (Fi~-6). The unusually large basal heat flow probably requires advective transport of heat within the crust (magma underplating or intrusion) or from an underlying abnormal mantle. Advective and conductive heat flows are not easily separated. High heat flows, up to 100-110 m W / m 2, are observed along

active margins (Watanabe et al., 1977; Henry and Pollack, 1988).

6. Deformation of the crust

The corresponding deformation map is compiled from temperature distribution with depth, a friction law (Carter and Tsenn, 1987) and a power law with the respective parameters (Table 1) assigned to each layer (Wilks and Carter, 1990). The frictional stress increases linearly with depth. The power law uses different values for the activation energy of granodiorite (Q¢ = 212 k J / m o l ) and amphibolites (Qc = 244 kJ/mol). The factor ruling the power law ranges from 2.4 to 3.7 as usually observed on crustal rocks. For a standard strain rate of 10-16 s - 1, the resulting

194

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

Table 1 List of the palameters used in the present model Layer Thickness Heat prod. (kin) (mW/m 3)

log(A) (MPa-" s

Granodiorite Amphibolitc Fetsic granulite

1.5 - 9.11 - 4.82

8 I0 13

2.9 1.7 0.2

temperature of the brittle/ductile zone is a little lower for granodiorite (about 220°C) than estimates from observations (300-350°C) on natural samples (Gapais, 1989). For amphibolites, the transition temperature (360°C) agrees better with the observations (400°C) (Tullis and Yund, 1991). Under such conditions, the brittle/ductile transition occurs at 8 km depth, corresponding to a temperature of 360°C and a stress level of 160 MPa for the onset of extensional fractures (Fig. 6). These are not threshold values, since brittle/ductile transition is strain, as well as fluid dependent. It also depends on the composition of the deformed rocks, and it is therefore affected by large-scale heterogeneities. Consequently, the transition zone ranges between 1 and 2 km from the proposed values. Anyway, the source area for granites is well within the ductile field, as expected from the elevated temperatures necessary to induce melting and most of the acid magmas' ascent takes place within the ductile field.

7. Brief review on segregation and collection of magma

After initial melting at grain boundaries, melt connection is established (Jurewicz and Watson, 1985) and magma can be collected. When enough melt establishes a continuous film at grain boundaries, it defines a first percolation threshold (Vigneresse et al., 1991) and magma starts to migrate. When the melt percentage increases to about 2 6 - 3 0 % , it defines a second percolation threshold (Van der Molen and Paterson, 1979; Vigneresse et al., 1991). The solid matrix breaks down and the melt plus the solid segregate escape fi'om the source area. Theoretical and experimental studies describe how melt segregates within a matrix (McKenzie, 1984; Ribe, 1987; Wickham, 1987), although they

I)

Qc (kJ tool I)

n

212 244 243

2.4 3.7 3.1

mainly concern basic magmas. Only a few theoretical studies relate to magma segregation and regional deformation in acid rocks (Wickham, 1987). Magma segregation is modeled by means of mathematical equations which incorporate the nonhydrostatic stresses existing within the system (McKenzie, 1984; Ribe, 1987). Buoyancy is the main driving force. Other forces such as shear and compaction of the matrix may act as driving or reacting forces. When the stress is low, magma upwelling is controlled only by the buoyancy of the melt. Magma ascent depends on the viscosity of the melt. Conversely, if nonhydrostatic stresses overcome buoyancy, then magma segregation is controlled by matrix viscosity (Ribe, 1987). Vertical forces such as suction or compaction have a positive, but minor influence on magma upwelling (Ribe, 1987; McKenzie, 1984). Regional deformation is such an external force which can segregate magma. In the case of shear, nonhydrostatic forces act as driving forces [right hand side in eq. 7b of Ribe (1987)]. Thus shear acts as a positive force to segregate magma. In plastically deforming materials, stresses vary according to their distance from the major area of deformation. They segregate magma up to the point that deformation can be enhanced if too much material with low viscosity is accumulated (Hollister and Crawford, 1986). Melt segregation has been described in small-scale experiments with very small amounts (less than 20%) of liquid (Dell'Angelo and Tullis, 1988; Cooper, 1990). In those experiments, the liquid percentage is not sufficient to model the behavior of a major granite pluton emplacement (about 70% of melt). However, even though large-scale transport is not realized, a small fraction of liquid clearly moves toward the extensional regions, as demonstrated experimentally (Cooper, 1990) and in field observations (Allison and Norris, 1992). Experiments also

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

emphasize the role played by matrix deformation which limits the rate of liquid migration, corroborating the numerical simulations (Ribe, 1987). To recapitulate, at a depth of incipient melting (17-20 km), liquid pockets initiate at grain boundaries and magma is collected when enough connections are made between pools, reflecting the increase in melt percentage (approx. 30%). At this stage, or second percolation threshold, the driving forces mainly result from the density and viscosity contrasts between melt and matrix, making diapiric ascent of magma possible because the crust is still ductile. However, the matrix must move fast enough to fill in the volume of the expelled magma (Ribe, 1987; Cooper, 1990), a process which is still more efficient if deformation is added to the hydrodynamic forces induced by the magma. Shearing appears to be a more efficient mechanism than compression (Ribe, 1987), which is also observed in migmatites (Dell'Angelo and Tullis, 1988; Allison and Norris, 1992).

8. Upwelling of magma in a ductile environment: channeling or diapirism?

Upwelling magma can be collected in veins or dykes through which a large amount of material flows. It has been described for basic magmas (Nicolas, 1986) as well as for acid ones (Clemens and Mawer, 1992; Petford et al., 1994). Magma is transported from its source toward the surface by discrete channels. Another model implies diapiric upwelling during which most of the melt and its matrix are transported to the surface. Diapiric processes have been proposed for salt domes and modeled for mafic magmas in the mantle (Marsh, 1979; Ramberg, 1981; Mahon et al., 1988; Singer et al., 1989). Magma is displaced because the density and viscosity contrasts from which Rayleigh-Taylor instabilities develop. Since structural data or the morphology of plutons reflect only the late stage of magma emplacement, they are not representative of its ascent. Field observations are also limited in this respect (Schwerdtner, 1990). However, the volume of the plutons may place constraints on the volume of the source areas. By nature, dykes imply a more or less continuous

195

flow into discrete conduits of minor size. The total volume of magma transferred cannot determine, neither the geometry, nor the number of dykes. But the viscosity of the magma and its flow rate prior to crystallization can limit the number of dykes (Petford et al., 1994). At those depths, gravity measurements have too low a resolution to determine the accurate size of the dykes. It depends mainly on the average spacing between measurements and accuracy. Average sample density is about 1 p t / k m 2. Therefore, bodies of smaller size cannot be identified, although their depth (about 5 kin) implies a long-wavelength component that widens the signal. Granitic bodies at that depth, about 100-200 m in radius, only create an anomaly of 0.1 mGal which is within the confidence level. This is well above the estimated width of dykes, about 7 m in Flamanville and 8 m in Mortagne (Petford et al., 1994). The alternative, diapiric upwelling, is associated with the onset of Rayleigh-Taylor instability (Ramberg, 1981; Weinberg and Schmeling, 1992). When a thick layer blankets a less viscous and less dense layer, the mobile layer soon becomes unstable and periodic domes develop, which may become detached from the source layer, giving way to buoyant diapirs. This is valid for a thin unstable layer beneath a thicker cover: h I > 10h 2

(1)

The spacing (A) of intrusive centers is ruled by (Selig, 1965): r/, ) '/3 A = 2.1527r h2\( r/---~2

(2)

Here, ~)j is the viscosity of the upper layer of thickness h l, overlaying a layer of thickness h 2 and viscosity "Oz. A supplementary equation calculates the volume (V) of the transferred magma relative to the volume of the source. In a tabular model of the source, the spacing between intrusions determines the volume of magma collection at depth. The total volume of the molten source must be greater than: v = ,~2h 2

(3)

This is certainly a minimum value, and would be multiplied by a factor 4 if one accepts that granites

J.L. Vigneresse / Tecmnophysics 249 (1995) 187-202

196

are derived from the melting of amphibolites in the bulk equation (Wedepohl et al., 1991): 4 amphibolite (a) --> 1 granite (3') + 3granulite(yv)

(4) 4V

=

A2h2

(5)

A combination of these equations is used (Fig. 7) to limit the thickness of the mobile layer as a function of the viscosity contrast. Viscosity of the overburden layer varies between 1020 and 10 t7 Pas depending on the ambient temperature. Estimated viscosities for molten magmas are about l0 s Pas

h2

x

6

o.1

2

k,~, 2

.s

~'~

..~

.ff%z/Z~>/

q

"qV" ,

V = X- h,

Y:7

/7/ -6

-3

0

Fig. 7. Parameters which control the diapiric upwelling of a volume V of magma from a layer with thickness h2 and viscosity •/2 (see text and figure in inset). Controls are: (a) the spacing (A) of plutons shown by thick lines with values (30-70 km); (b) the volume of the molten layer, shown in large stippled lines if the whole layer melts, or in small stippled lines to account for a 25% of partial melting. The resulting unshadowed box corresponds to the range of parameters (viscosity, thickness) in case of diapirism. The dotted zone on the left indicates the thickness of the molten layer which can produce the required volume of magma with the corresponding spacing of plutons (600-2000 km).

(Rubin, 1993) ranging between 10 4 and 108 Pas for rhyolite (Clemens and Mawer, 1992). According to Eq. 2, the average spacing between intrusions is a direct function of both the viscosity contrast and the thickness of the buoyant layer. Spacing between plutons varies from about 32-35 km for plutons and volcanoes in the Lachlan told belt (Rickard and Ward, 1981) to about 70-80 km for plutons in island arcs (Vogt, 1974; Marsh, 1979). The latter value is observed in southern Brittany (70 km). A pluton interval ranging from 30 to 70 km implies a thin unstable layer, a kilometer or less thick (see also Marsh, 1979). It also implies viscosity contrasts of about 10 3 to 10 - 6 , yielding values of 101! to 1014 Pas for the unstable layer, which is far above the experimental data. The mobile layer must be ten times thinner than the overlying layer (Eq. 1). If the average crustal thickness is overestimated by a factor 2, 60 km thick, the mobile layer cannot exceed 6 kin. The last limitation comes from the volume equation (Eq. 3). When all source material is transferred to intrusions, this volume provides the minimal thickness of the source layer. When Eq. 5 is applied, which relates granitic magma to partial melting of amphibolites, it is clear that not all three equations (Eq. 1, Eqs. 2 and 5) be simultaneously fulfilled (Fig. 7). These arguments are not in favour of diapiric upwelling in the lower crust. Alternatively, when adequate viscosity contrasts are applied (at least 1 0 - 6 ) , the observed spacing implies an unstable layer 1-24 m thick. The volume of magma generated by that layer is one or two orders of magnitude less than the volume of intruded magma. Conversely, to achieve magma volume conditions, the viscosity contrast (10 -3) does not match those measured in molten magmas. To produce the required volume of acid magma (about 1500 km3), the molten part of the crust must be 0.3 km thick, if the spacing between plutons is approx. 70 kin. This implies that at least 1.2 km of amphibolites is to melt. A thickness of 2 km is proposed since not all melt can be extracted and transferred to plutons. This also corresponds with the approximate melting percentage (26-30%) above which the partly molten mixture loses cohesion and allows magma to segregate (Van der Molen and Paterson, 1979). The wavelength of diapirs developed by this unstable layer, assuming a viscosity

197

J.L Vigneresse / Tectonophysics 249 (1995) 187-202

I

200 I

I

,,

400 I

600 I

I

,,

\

I ~ -8°° MPa

I

800 I

I

I

T °C

\

\\9

~

~ ~

~.

Iting

~/~ !.~~increasing basalq

Fig. 8. Pressure-temperature diagram of granite formation and &scent in a three-layered (granodiorite/amphibolite/granulite) crust. Conductive geotherms are indicated with dashes, with values of their basal heat flow (Q). Initial melting region is shown by large dots. Ascent of the granitic melt is mainly within the ductile field and follows an adiabatic curve, then heat transfer takes the path toward lower temperatures. This cannot be included in the diagram because advective heat transfer distorts the temperature distribution with depth. Crystallization (in grey) occurs after the brittle/ductile transition zone (dotted) is reached. WS = wet solidus; WP = wet solidus for pelites; DP = dry pelites solidus; Am = melting curve for amphibolites.

contrast of 10 -6 with the host rocks, is 684 km, one order o f magnitude greater than the one observed. These findings suggest that diapirism is not an efficient mechanism to feed acid magma, therefore dykes are proposed as an alternative (Clemens and Mawer, 1992; Rubin, 1993; Petford et al., 1994). Consequently, the ascent o f granitic m a g m a is explained as follows (Fig. 8). The first part o f the ascent is close to the adiabatic curve, melting is not enhanced during the ascent, but the heat carried by the m a g m a is insufficient to melt the crust. That heat loss directs the ascent o f the m a g m a towards lower temperatures, reaching the b r i t t l e / d u c t i l e transition at temperatures o f ca. 350-400°C. Ascent is achieved through small dykes, although the exact size and number cannot be estimated from the information at hand, but a range in the tens o f meters is suggested (Petford et al., 1994). In spite o f extensive research, no relationship has ever been established between a pluton and its source. The feeder channels remain

unseen. Numerous dykes present under the Himalayan granites (Le Fort et al., 1987) support the dyke hypothesis rather than diapirism. Small rounded outcrops o f both acid and basic magmas in the Smartville Complex (Beard and Day, 1987) have been proposed to represent the eroded feeder channels of that pluton.

9. T h e ascent o f m a g m a f r o m a ductile to a brittle crust

The depth o f emplacement of the granite o f Pontivy (Brittany) has been estimated from the index minerals of its metamorphic aureole. Lack of kyanite in the mineral assemblage implies a pressure o f approx. 200 MPa, whereas illite crystallinity provides temperature estimates o f approx. 2 0 0 - 2 5 0 ° C (Le Corre, 1975): Using those numbers, the granitic roots, at approx. 4 - 5 km beneath the level o f the

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metamorphic aureole, were not far above the b r i t t l e / d u c t i l e transition zone (Fig. 9). The major barrier faced by the m a g m a during its ascent is the crossing of the transition zone from a ductile to a brittle crust. According to the above thermal model, this occurs at approx. 400°C for a feldspar-rich crust and at approx. 300-350°C for a quartzitic crust (Gapais, 1989; Tullis and Yund, 1991). The temperature of transition is independent of the rate o f strain and has little effect on the strain level required to fracture the crust. In both types of material, some extra stress is needed to initiate fractures within the brittle crust, approx. 150-200 MPa for extensional fractures. In case o f strike slip or thrust faulting, these values must be multiplied by a factor 2 to 4. The hydrodynamic forces of an ascending column of m a g m a are unable to develop such a stress. The density contrast between the m a g m a (2400 k g / m 3) and surrounding crust (2750 k g / m 3) is low, approx. 350 k g / m 3. The hydrodynamic force caused

by the lighter specific gravity of the liquid is approx. 3.5 MPa per 1 km height of magma. That value is two orders of magnitude less than the stress value required to penetrate the b r i t t l e / d u c t i l e zone at a depth of approx. 10 kin. Incorporating a fluid phase into the magma lowers its density, and thus increases the hydrodynamic force. However, even if 25% of fluids is added, it only doubles the density contrast as well as the hydrodynamic force, but the latter still fails to overcome the required stress value to intrude the upper brittle crust. Therefore an unrealistically tall column of granitic m a g m a is required to overcome the shear resistance of the crust diapirically. Nevertheless, small dykes or conduits can develop because their propagation is catastrophic due to the larger forces developed at their tip (Lister and Kerr, 1991; Clemens and Mawer, 1992). In the case of the Flamanville pluton, the very small volume of m a g m a calculated from gravity data ( B r u n e t al., 1990) is definitively insufficient to enable diapiric transport

10 km 10_ 20 i

b!salheatflow

Segregation and ascent

collect and emplacement

Fig. 9. Model of granite emplacement according to a realistic crustal model. Vertical and horizontal scales are identical. The brittle/ductile transition zone (dotted) is calculated using an average thermal gradient of approx. 37°C/km. Melting (large dots) starts at 700°C and 600 MPa according to geochemical data. Magma wells up to the upper crust and may pool near the transition zone from ductile to brittle deformation. The shape of the granitic body is taken from Pontivy, according to gravity data. The eroded part of the pluton has been extrapolated. Depth of emplacement is determined from the metamorphic aureole mineral assemblage. In that case, the roots are very close to the brittle/ductile transition zone.

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

to overcome the brittle/ductile stress level and a dyke of approx. 7 m in width has been suggested instead (Petford et al., 1994). Structural arguments also favour upwelling mechanisms which involve other forces, mostly induced by deformation, than magma buoyancy and viscosity contrast, which operate during diapirism. In the Flamanville pluton (Brun et al., 1990), asymmetrical deformation of the surrounding sediments, magmatic shear zones within the granite, and a very small conduit ( < 300 m from gravity, tens of meters according to the model) at a shallow depth (3 km) oriented along the axial fold plane of the surrounding deformed sediments, suggest that the competence of the surrounding rocks has controlled the emplacement of the granitic body. In Mortagne and Cabeza de Araya, fractures opened deep enough to involve the ductile crust, thus providing space for the magma to penetrate the upper brittle crust. The strength required to break the brittle crust is about four times greater under compression than under extension. Consequently, the easiest place for magma to intrude the crust in a mechanical sense, is within the extensional areas where the stress level is the lowest. It does not imply that granites can only intrude extensional areas. Extension always exists locally as a proximal effect, even in a regional compressional stress field. The peculiar emplacement of granites within local extensional areas is currently observed in Brittany, in Galicia, or in plutons emplaced during transtension (Guineberteau et al., 1987; Tikoff and Teyssier, 1992). According to the geometry calculated from gravity data, most granitic plutons are relatively thin, with a vertical dimension far less than their lateral extent. That shape differs from models observed during experimental diapirism, which look like an inverted tear drop. However, the argument is insufficient to disprove the diapiric mechanism of upwelling, since the flat shape could result from lateral spreading after the ascent of magma. But, when all parameters are considered, such as the thin shape of the plutons, the viscosity and volume involved, the low stress created by the magma itself, diapirism should be restricted only to regions which deform plastically. Additional evidence for diapiric upwelling requires structural observations which clearly demonstrate the vertical displacement of near-by

199

country rocks (England, 1990), radial structures or large folds, basins and denser rocks at the contact with the diapir (Schwerdtner, 1990), but that is not always the case. When approaching the brittle/ductile transition, the quantity of magma delivered by each dyke may be pooled and turned into sills, because of the change in stress level (Rubin, 1993). Some of these magma pools may feed the upper crust, because deformation of the overlying brittle crust also determines the level of pooling (Mortagne, Cabeza de Araya). During the ascent and at the stage of pooling, melting is certainly complete or sufficient to allow for convection and homogeneisation of magma.

10. Localisation of granitic intrusions Granitic plutons are often observed distributed over a large area as small intrusions associated with large-scale structures. For instance, in southern Brittany, four to five granites crop out which are associated with a large shear zone running trough the area. The coastal batholith of Peru is also a clear example of plutons linked by a linear structure along thousands of kilometers (Pitcher, 1979). It is, therefore, a popular idea to link granite genesis, ascent and emplacement with these structures. I have already discussed the possibility of shear melting (Vigneresse, 1995). In my opinion, this process cannot produce sufficient acid melt to generate a pluton. In other words, to produce 1500 km 3 of magma, with a spacing of 70 km along a fault extending down to approx. 20 kin, implies complete melting along a 1.1-kin-wide zone. The degree of melting directly controls the width of the softened zone. However, a large melt percentage makes rocks lose their internal cohesion and enhances deformation (Van der Molen and Paterson, 1979; Hollister and Crawford, 1986). A limit of 26% melt implies a 5 km wide fault and it seriously lowers the stress level in the vicinity of the fault. Observations on heat flow and stress levels near the San Andreas fault, from the Cajon Pass deep drilling project, do not meet this requirement (Lachenbruch and Sass, 1992). Such a width is also in contradiction with the concept of shear melting by which the melt is induced by the energy dissipated in the movement. If the

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shear width increases, energy dissipates more and melting is less important. Therefore, melt is not induced by deformation. Ascent of magma within fault zones is a postulated mechanism. To achieve spatial concentration of magma, its flow must have a horizontal component in addition to the vertical one. Lateral movement of magma is plausible in transcurrent shear zones, because of the horizontal movement along the shear planes. However, the transition from a melt film to a pluton bordered by shear planes, implies that magma has been halted in its lateral movement by some kind of blocking mechanism. This occurs if geometrical irregularities exist, relating to transtension or transpression (Tikoff and Teyssier, 1992). In that case, magma transport is certainly induced by the movements along the fault. The space created between overlapping branches of the fault localizes magma emplacement (see Mortagne, for instance). In all other cases, that is when the fault has no extensional component, the ascent of magma cannot be related to the fault zone. One conclusive observation is the systematic offset of the root zone from the major shear plane (Vigneresse, 1983, 1995). According to the arguments developed before, diapirism is not an appropriate mechanism to control magma ascent. Nevertheless, the spacing of plutons remains to be explained. The present paper, as well as its companion paper (Vigneresse, 1995), both locus on the role of deformation during the segregation, ascent and emplacement of acid plutons. Deformation creates local zones of extension in which pluton emplacement is facilitated. If extensional zones are periodic, then pluton distribution shares the same wavelength as extension. Periodic instabilities develop because of the non-linearity in the stressstrain profile of the brittle and ductile parts of the crust and underlying mantle (Martinod and Davy, 1992). Associated wavelengths develop, approx, five to six times that of the plastic-viscous boundary in the mantle, and about five times the thickness of the brittle crust ( 4 0 - 6 0 kin). For a thermally perturbed system, resulting in a hotter mantle, losing its plastic character, the large-scale wavelength disappears and buckling or boudinage are observed during compression or extension (Martinod and Davy, 1992). Those values ( 4 0 - 6 0 km) are in the range of those corresponding to the spacing of plutons. Thus, pluton

spacing should not be related to periodic diapiric instabilities, but to periodic extensional areas in a brittle/plastic deforming crust.

11. Conclusions Structural and gravity data help to locate root zones within granitic plutons. Their number and location are controlled by regional deformation, thus emphasizing the role played by regional deformation during the emplacement of magma. The volume of magma issued from one conduit is about 1500 km 3. The resulting shape of the pluton shows a horizontal/vertical ratio of about 10. Unlbrtunately, those data which reflect the final stage of magma emplacement do not explain the ascent mechanism. A thermal model of the crust has been constructed using the present structure and heat generation within the Hercynian crust. To achieve melting in order to produce granites, a large amount of heat must be supplied. It is proposed that deformation is a necessary condition for granitic magma to segregate, to well up and be emplaced in the crust. Lower in the ductile crust, deformation contributes to segregation and ascent. Diapirism appears to be restricted to very specific conditions and is not pertinent to the upwelling of granitic plutons. The volume of the produced magma requires a thick source layer which in turn implies a reduced spacing between plutons, which is not observed. The dyke hypothesis is therefore favored which also agrees more with the final morphology of the pluton. When reaching the brittle/ductile transition zone, magma is unable to overcome the stress level on its own and may start to pool. There, additional stress is required to break the upper brittle crust. The field of deformation also induces anisotropic (extensional or compressional) regions in the crust into which magma can enter. In that sense granites should all be synkinematic (Karlstrom, 1989). Deformation is not the cause of melting, nor localizes upwelling of magma except in the transtensional or transpressional areas. Granitic plutons are a consequence of deformation in the sense that segregation and ascent of magma are controlled by the field of deformation.

J.L. Vigneresse / Tectonophysics 249 (1995) 187-202

Acknowledgements This p a p e r is the result o f m a n y field s u r v e y s c o n d u c t e d o v e r granitic areas with the h e l p o f stud e n t s o f the U n i v e r s i t i e s o f B i l b a o , M o n t p e l l i e r , M u n c h e n , N a n c y , N a n t e s , R e n n e s , S a l a m a n c a and T o u l o u s e . C o n s t r u c t i v e d i s c u s s i o n s with structural g e o l o g i s t s , a m o n g s t o t h e r s , J.L. B o u c h e z , K.J.W. M c C a f f r e y and W . M . S c h w e r d t n e r , h a v e e n h a n c e d the r e a s o n i n g and i m p r o v e d the m a n u s c r i p t . I specially a c k n o w l e d g e J i m B e a r d for d i r e c t i n g m y attention to the S m a r t v i l l e C o m p l e x , C A , w h e r e s u p p o s edly e r o d e d r o o t s can b e o b s e r v e d . F u n d i n g for the gravity and t h e r m a l s u r v e y s has b e e n r e a l i z e d t h r o u g h grants

from

the

CNRS,

INSU,

DBT

program,

C R E G U and f r o m the m i n i n g c o m p a n i e s C o g e m a and Cisa.

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