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39Ar and RbSr analyses from ductile shear zones from the Atacama Fault Zone, northern Chile: the age of deformation
ELSEVIER
Tectonophysics 250 (1995) 61-87
JOAr/ “Ar and Rb-Sr analyses from ductile shear zones from the Atacama Fault Zone, northern Chile: the age of deformation Ekkehard Scheuber a,*, Konrad Hammerschmidt
b, Hans Friedrichsen
b
’ lnstitutfir
Geologic. Geophysik und Geoinformatik, Freie Uniuersitiit Berlin, Malteserstrasse 74.100, D-12249 Berlin, Germany h Institutfir Mineralogie, FR Geochemie, Freie Universittit Berlin, Boltzmannstrasse 18-20, D-14195 Berlin, Germany
Received 19 July 1994; accepted 6 April 1995
Abstract The influence of deformation on the K-Ar and the Rb-Sr isotope system is investigated. It is assumed that, due to the diffusion processes involved, deformation has a similar effect on isotopic equilibrium as has temperature. In order to examine the influence of deformation on the K-Ar and the Rb-Sr isotope systems two shear zones from the Atacama Fault Zone (AFZ), situated in the north Chilean Coastal Cordillera, have been investigated. The AFZ, which was active as a sinistral strike-slip fault during the Mesozoic, has two sets of shear zones, one formed under amphibolite (SZl), one under greenschist facies conditions (SZ2). Rb-Sr and 40Ar/39Ar age determinations were conducted on samples from cross sections of each set. In SZl the homblendes and biotites from a weakly deformed sample reveal cooling ages of 153-152 and 150 f 1 Ma, respectively. Biotite from the center of the shear zone of SZl gave an isochron of 143.9 + 0.3 Ma (MSWD = 0.04) which is interpreted as the age of deformation which produced resetting of the mineral system. In SZ2 homblendes yielded “OAr/ 39Ar plateau (cooling) ages of - 138 Ma. Biotites from undeformed samples gave Rb-Sr and “OAr/ 39Ar total degassing ages of 130 f 1 Ma, whereas biotite from the mylonitic rocks yielded 126- 125 Ma which dates the time of deformation. Sr isotope homogenization occurred in the mylonitic rocks, and is most likely a result of deformation. The formation of SZl can be correlated to the Araucanian (= Nevadan) phase. The deformation in SZ2 is related to the onset of uplift and cooling of the Coastal Cordilleran magmatic arc.
1. Introduction Geochronological methods are a powerful tool for the dating of semiductile and ductile deformational structures (e.g., Kligfield et al., 1986; Hubbard and Harrison, 1989; Treloar et al., 1989; Getty and Gromet, 1992; Leloup et al., 1993). However, dating deformational events often meets the problem of how to distinguish ages due to regional cooling from
* Corresponding author. 004O- 195 l/95/$09.50 0 1995 Elsevier Science B.V. All rights reserved LSD1 0040- I95 1(95)00044-5
ages due to deformation because ductile deformation of quartz-feldspar rocks is only possible in nature under elevated temperatures (greenschist facies or higher). Cooling ages are associated with the closing of isotopic systems of a mineral by passing through its closure temperature CT,>. In the relation between the age of a mineral, recrystallized during deformation, and passing the T, two cases can be distinguished: (1) When deformation takes place before the mineral falls short of T,, the age of this mineral is a pure cooling age which only gives a minimum age of deformation. (2) When deformation takes place after passing T,, it is possible to determine the
62
E. Scheuber et al./ Tectonophysics 250 (1995161-87
age of deformational events if the mineral under question recrystallizes and attains complete isotopic equilibrium. In literature it is assumed that the critical role in age resetting is played by mass transfer due to diffusion (Dodson, 1979). Diffusion is facilitated not only by elevated temperatures, but also by differential stresses and strain energy (e.g., Wintsch and Dunning, 1985; Wheeler, 1987) which have a great influence on chemical potentials and, for example, are driving forces for metamorphic reactions (Wintsch and Andrews, 1988; Klaper, 19901. In this paper we discuss the influence of deformation on the Rb-Sr and K-Ar systems of mylonites from the Atacama Fault Zone (AFZ), northern Chile; here, two sets of ductile shear zones occur which originated under amphibolite facies and greenschist facies conditions, respectively (Fig. 1). Age determinations of biotite and hornblende have been carried out in cross sections from the undeformed protolith into the center of shear zones crosscutting plutonic rocks. It will be shown that deformation produced a resetting of Rb-Sr and K-Ar mineral ages, and that this resetting was accompanied by homogenization of the Rb-Sr whole-rock system at least on the meter scale. However, resetting and homogenization was only detectable in minerals of sheared rocks where deformation started after cooling of the undeformed protolith below T,; the age of high-grade deformations in an amphibolite facies shear zone gave only the age of cooling after deformation. I .I. Plastic deformation
and closure temperatures
The theory of closure temperatures (Dodson, 1973, 1976, 1979) is based on the assumption of a firstorder loss or of volume diffusion of the radiogenic nuclides (e.g., Chermiak and Watson, 1992). T,, which is defined as the temperature (T,) of the system at its calculated age, is given by: E DO = In (~72 i i RT,
Fig. 1. Map of the north Chilean Coastal Cordillera showing the location of shear zones along the Atacama Fault Zone, and the investieated cross sections SZl and SZ2.
E. Scheuber et al./
Tectonophysics 250 (1995) 61-87
where E is the activation energy, R the gas constant (R = L X k, L is the Avogadro constant, k is the Boltzmann constant), D, the frequency factor (diffusion coefficient at infinitesimally high temperature), a the radius of the mineral grain, (Y a geometric factor and r a cooling-time constant (r = - RTcz/[ E(dT/dt)]} which considers the cooling rate (dT/dt) of a particular area (T is temperature and r is time). In plastic deformation of silicate rocks under natural conditions the most common deformation processes are diffusional flow and dislocation creep. In both processes the strain rate is controlled by the applied stress and by the diffusion of vacancies, ions or atoms (e.g., Poirier, 1985). Stress (a) and strain rate (+) are related by the linear form (cf. Frost and Ashby, 1982; Poirier, 1985): D eff in the case of diffusion form:
creep, and by the power-law
&=A( gneXP( - $,) in the case of dislocation creep (power-law climbplus-glide sensu Frost and Ashby, 19821, where (Y is a numerical factor depending on the shape of the grain, R is the atomic or molecular volume, T is the temperature, a is the grain size, De,, is the effective diffusion coefficient, A is the power-law creep constant, p is the shear modulus and n is the stress exponent (stress sensitivity). Dislocation creep causes lattice defects to migrate through crystals under an applied stress to form suband new-grain boundaries by either migration recrystallization or rotation recrystallization (cf. Drury and Urai, 1990). Migration recrystallization is in particular expected to support chemical and/or isotopic exchange (Getty and Gromet, 1992), as in this process regions of high dislocation density are repeatedly replaced by dislocation-free material as mobile grain boundaries migrate through crystals. Diffusion creep is the predominant deformation process at very small grain sizes ( < 10 pm for quartz, Handy, 1989); it is also expected to support chemical and isotopic exchange as it results from stress-directed
63
bulk vacancy diffusion (Tsenn and Carter, 1987) along grain boundaries or through the lattice. Diffusion creep includes pressure solution which is strain and stress controlled. For example, Wintsch and Dunning (1985) have shown that the solubility of quartz increases if tangled dislocations are present (‘‘strain solution’ ‘). Grain-scale differential stresses induce pressure solution and are a driving force for chemical exchange (Wheeler, 1987; Wintsch and Andrews, 1988). Although at present no theory is available which links the theories of closure temperature and plastic deformation, a comparison of both suggests that, due to the diffusion processes involved, both strain rate and temperature have a similar effect on closure or opening of isotope systems. This means that an increasing strain rate should lead to age resetting in an environment of uniform temperature, even below T, and although resetting may be enhanced by shear heating. Therefore, if a gradient in strain rate is observed, age resetting should preferably affect those rocks which accommodate higher strain rates. This means that, in inhomogeneous materials where the strain rate concentrates on the weaker parts, age resetting should preferably affect the weaker material. Furthermore, if the strain rate drop after deformation is rapid, the relevant isotope systems of all minerals that recrystallized dynamically will be “frozen in”. This is similar to the preservation of microstructures which also depends on the rate of strain rate drop (Knipe, 1989). On the other hand, a slow decrease in stress/strain rate should have the same effect as if the temperature remains high: the deformational microstructures should be obliterated by annealing; in this case the minerals are expected to show a normal cooling-like pattern, and the minerals show different ages. The finite strain is also expected to support age resetting as, with increasing finite strain, the number of grains undergoing dynamic recrystallization increases. This is similar to the time dependence of heating on isotopic exchange: the duration of heating controls the degree of isotopic exchange. A very important effect of dynamic recrystallization which is expected to support isotopic exchange is the deformation-induced grain size reduction. The grain size-stress relation (e.g., Twiss, 19771, which has been established theoretically for migration re-
64
E. Scheuber et al./ Tectonophysics
crystallization by Derby (1990), says that the smaller the recrystallized grain size the higher the applied stress. Smaller grain sizes, on the other hand, reduce the closure temperature (Eq. 1). Therefore, it is possible to observe large but older grains whose age values survive in a matrix of younger but smaller grains. From these considerations it is concluded that the age of a deformational event can be dated if strain rate, strain and/or deformation temperature are sufficient to produce a resetting of the investigated isotope systems. To distinguish an age due to deformation from an age due to regional cooling the following two criteria should be fulfilled: (1) deformation should have started below the closure temperature of the host rock; and (2) deformation-induced resetting should be complete. Criterion (1) is fulfilled if the same mineral species is younger in the tectonite than in the host rock. Criterion (2) is fulfilled if deformed or recrystallized minerals show isotopic equilibrium with other minerals.
2. Geological
setting
During the Jurassic and Early Cretaceous a continental magmatic arc developed in the north Chilean Coastal Cordillera. It consists of large quantities of basaltic to andesitic lavas (La Negra Formation, Garcia, 1967) and extended batholiths of gabbroic to granodioritic composition. Magmatic activity lasted from the Early Jurassic to the Early Cretaceous, as shown by isotope age data (200-I 10 Ma; e.g., Halpem, 1978; Diaz et al., 1985; Maksaev et al., 1988; HervC and Marinovic, 1989). The tectonic setting of the Jurassic-Early Cretaceous arc can be described as extensional to transtensional (Scheuber and Reutter, 1992; Grocott et al., 1994). Its major structure is the orogen-parallel Atacama Fault Zone (AFZ) which can be traced for more than a thousand kilometers from _ 20” to N 30%. Scheuber and Andriessen (1990) gave a detailed description of AFZ structures south of Antofagasta (23”45’-25%) where two sets of ductile shear zones developed (Fig. l), one formed under amphibolite facies conditions and one under greenschist facies conditions (called Jurassic and Early
250 (1995) 61-87
Cretaceous shear zones, respectively, by Scheuber and Andriessen, 1990). In the Early Cretaceous shear zones the vertical foliation planes generally trend N20”E and contain subhorizontal stretching lineations. The sense of shear is uniformly sinistral, which is consistent with highly oblique ( > 45”) subduction towards te southeast during the Late Jurassic-Early Cretaceous (Zonenshayn et al., 1984, Jaillard et al., 1990); the AFZ of that time was thus interpreted as a trench-linked strike-slip fault sensu Woodcock (1986). In the Jurassic shear zones the general trend of the foliation planes is also N20”E. However, kinematic indicators are less well developed and rarer than in the Early Cretaceous shear zones probably due to thermal or high-temperature tectonic overprinting. The kinematic pattern is complicated and not yet fully understood. Structures indicative of sinistral strike slip were overprinted and partly obscured by vertical displacements in some places (Gonzalez, 1993). Scheuber and Andriessen (1990) have published evidence for sinistral strike-slip movements. Structures due to vertical displacements have been described by Brown et al. (1993) from the Coastal Cordillera between 25 and 27%. These authors attribute such movements to the subsidence of a marine backarc basin to the east of the magmatic arc. In some places an overprinting of the structures due to strike-slip movements by dip-slip normal movements can be observed. One example is illustrated in Fig. 2. At the southern contact of the Pluton Cristales (Rb-Sr whole-rock age 145 + 10 Ma, Sr, 0.70326, Her& and Marinovic, 1989) the overall N-S trend of the foliation planes is overprinted by a second generation of planes running parallel to the contact of the Pluton Cristales. Near the contact the older planes are completely obliterated, whereas at a greater distance from the contact both planes can be distinguished. The older foliations contain subhorizontal stretching lineations, and are thus due to strike-slip movements; on the other hand, the younger foliations contain down-dip lineations, and kinematic indicators reveal subsidence of the Pluton Cristales with respect to its host rock. The older (strike-slip) deformation led to the formation of a recrystallized hornblende-plagioclase fabric, whereas the younger foliation planes are defined by the preferred orientation of biotite.
65
E. Scheuber et al. / Tectonophysics 250 (I 995) 61-87 70°25’
70030’ w
24’=05’
v;,x\,x x v \\ x--x-
x /_
,”
x x
-<, a{v
-iI
vv’
v v v
_&& ly
-
v-7
.‘\
>
a
“I
vvvvv vvvvv vvvvvv vvvvvv
I
24OlO’
v7
, 1’1
0
,‘I’,
m c
I’:‘ \ ?\ x 6,
v V - 24’=15’
5 km 1
2
3
4
2. Geological map of the Coastal Cordillera near Cerro Cristales showing the location of the investigated shear zone SZI. I = Jurassic intrusive rocks (Paranal Unit, - 170 Ma); 2 = Jurassic volcanic rocks; 3 = Pluton Cristales (- 143 Ma); 4 = trend of foliation planes; sources: own mapping and Gonzalez (1990).
Fig.
3. Lithology and structures 3.1. Amphibolite
facies shear zone (SZI)
SZl, approximately 15 m wide, is located N 50 m south of the contact to the Pluton Cristales (Figs. 1 and 2). Table 1 shows the locations, modal compositions, and structural features of the investigated samples. The vertical foliation planes of SZl strike N60”E, the older planes cannot be observed macroscopically. Due to increasing strain, the spacing of
foliation planes becomes narrower approaching the shear zone center. The shear zone is developed in a hornblende gabbro of the Paranal Unit (Rb-Sr whole-rock age 170 f 28 Ma, Sr, 0.70315, HervC and Marinovic, 1989) containing some biotite, orthoand clinopyroxene (sample 87-150). In the shear zone the gabbro is altered to an amphibolite gneiss (samples 87-153 10 m north of the shear zone center, 87-151 shear zone center, Fig. 3) consisting mainly of plagioclase and magnesio-hornblende. The Ancontent of plagioclase decreases with increasing strain
66
E. Scheuber et al. / Tectonophysics
from N 45 mole% (87-153) to 32 mole% (87-151) (microprobe analyses). In the shear zone center some newly crystallized orthopyroxene is present (Fig. 3a) indicating upper amphibolite facies metamorphic conditions (Bucher and Frey, 1994). The biotite content increases from N 1% in the undeformed gabbro to 7-9% in the shear zone probably as a result of a better permeability for fluids due to the smaller grain size in the shear zone. The growth relationships of biotite reflect both deformations which affected the SZl rocks; the following types of biotite grains can be distinguished: (1) old grains of magmatic origin; (2) flakes which have grown across the recrystallized plagioclase-hornblende fabric without preferred orientation, and which are frequently bent (Fig. 3b);
250 (I 995) 61-87
and (3) partly recrystallized grains with straight grain boundaries oriented parallel to the foliation (Fig. 3~). Old grains (1) form only a minor fraction of biotite, their occurrence being restricted to the undeformed (87-150) and the moderately deformed samples (87153); grains of type (2) also occur mainly in the moderately deformed sample, to a lesser extent also in the shear zone center; biotite grains of type (3) dominate in the shear zone center but they are also present in the shear zone margin. From the growth relationship between biotite, plagioclase and homblende the following run of events can be inferred: biotite of type (2) grew statically after the first deformation which produced the recrystallized plagioclase-hornblende fabric. During the second defor-
Table 1 Location, modal composition, and structural features of the analyzed samples: Modal composition (vol.%) of rock-forming minerals, the amount of recrystallized (recr.) and magmatic (magm.) grains; average grain size of matrix grains; and ellipticity of the finite strain ellipse in X2 sections [ R(XZ)]. In SZl amd SZ2, the samples are grouped from left to rigth, respectively, with decreasing distance from the shear zone center
SZl 87-150 70025.5’W 2.4011.5'S
SZ2
87-153
87.151
70'25.5'W 24011.5'S
70025.5'W 24or1.5s
85-58 70*15.5'w !4040.
I’S
88-26
88-1 a
70~17.3'W 24o38.1'S
70'19.3'W 24O37.I'S
88.lc 70~1.9.5'W 24O37.6'S
88-10
88-16
85.546
70~1.9.8'W 24O37.3'S
70"20.6'W 24038.3'S
70°20.6% 24O38.3'S
1.o
7.5 0.5 8.0
0.0 0.0 0.0
0.0 0.0 0.0
21.6 0.0 21.6
il.4 0.0 21.5
17.5 6.2 23.7
4.3 20.5 24.8
18.6 19.6
0.1 30.3 30.4
0.0 13.8 13.8
56.9 6.5 63.4
35.6 27.9 63.7
3.5 61.7 65.2
53.3 0.0 53.3
58.4 0.0 58.4
51.4 0.0 51.4
46.4 1.7 48.0
47.9 3.3 51.3
41.4 3.3 44.4
0.9 73.2 74.1
0.8 0.0 6.6
0.9 0.0 0.9
0.0 0.0 0.0
10.7 0.0 10.7
6.6 0.0 6.6
12.6 0.0 12.6
8.4 3.0 11.4
5.2 1.1 6.4
9.7 6.1 15.7
0.0 0.0 0.0
23.7 0.0 23.7
9.2 14.0 22.2
0.2 22.9 23.1
5.9 0.0 5.9
6.8 0.0 6.8
5.5 0.0 5.5
5.9 0.5 6.4
5.2 1.1 6.3
1.6 0.1 1.7
0.0 0.0 0.0
1.2 0.0 1.2
3.3 5.6 8.9
0.2 6.6 6.8
7.1 0.0 7.1
5.6 0.0 5.8
6.0 0.0 6.0
2.5 6.7 9.2
2.7 11.6 14.3
1.6 4.7 6.3
0.0 0.0 0.0
1.7at
0.0
o.1")
0.0
0.0
0.0
0.0
0.0
0.0
lZ.lC'
1.2
4.3
4.8
1.0
0.8
0.7
1.5
1.5
0.0
0.120
0.046
0.044
0.045
0.004
2.3
3
3.5
0.320
0.135
0.105
1.4 _.. -1
a Igneous pyroxenes. b Metamorphic pyroxenes. ’ Porphyroclasts.
._. -1
-1
>>50
61
E. Scheuber et al. / Tectonophysics 250 (1995) 61-87
The rocks of SZl were deformed within a fully ductile rheology without any ductility contrast between each mineral species. This is revealed by a steady increase in the amount of recrystallized grains of all mineral species from the margin into the center of the shear zone (Table 1; Fig. 41, and by the observation that there is no grain size reduction around old grains which thus did not operate as rigid particles but rather are remainders due to incomplete recrystallization. In the shear zone center recrystallization is nearly complete, whereas in the sample from the margin 40% of old grains are present. Postdeformational annealing in the rocks of SZl is revealed by a granoblastic fabric of the matrix where many grains have straight grain boundaries and 120” triple points are frequently found in plagioclase domains. These observations indicate that the peakstress event is not preserved in the microstructures of the rocks (Knipe, 1989; Prior et al., 1990) due to a relatively slow regional cooling and/or a slow decrease of stress after deformation. In the annealed fabric of the rocks from SZl it is not possible to determine the finite strain. 3.2. Greenschist facies shear zone (SZ2) SZ2 is located in a shear zone which runs through a medium-grained granodiorite of the Remiendos Unit (Her& and Marinovic, 1989: Rb-Sr biotite age
l-
Fig. 3. Microphotographs of sheared rocks of SZI. (a) Newly (Opx) in sample (87-1.51) from the crystallized orthopyroxene shear zone center. (b) Biotite flakes (B) which grew across the recrystallized, pophyroclastic fabric of plagioclase and hornblende (sample 87-153, 10 m north of shear zone center). (c) in the shear zone center (sample 87-151) biotite is recrystallized dynamically.
g
0.8 . .
f 8 2
0.6 .a
g r h g
?? 0.4 . . 0.2 a* 0*.. 0
A :
0.2
.
: 0.4
.
: 0.6
AA : 0.8
1
(qtz+bi)recr/qtz+bi)tot mation these grains were deformed and recrystallized to form type (3). Thus, recrystallized biotite (type 3) should be younger than grains of type (21, and of course also younger than igneous biotite of type (1).
Fig. 4. Diagram showing the degree of recrystallization (recr = recrystallized: tot = whole rock) of incompetent minerals (qtz = quartz; bi = biotite) and competent minerals (fsp = feldspar; hb = hornblende); squares: SZl, triangles: SZ2.
68
E. Scheuber et al./ Tecronophysics 2.50 (I9951 61-87
13 1 i- 1 Ma). The exposed width of the mylonitic zone is about 2 km. Samples have been collected from the undeformed granodiorite (samples 85-58 and 88-26), and across a section through the shear zone (samples 88-la, c, e, g and 8554b) (Table 1; Fig. 2). A NNE-striking subvertical foliation develops N 1.9 km east of the shear zone center. Due to increasing finite strain the spacing of the foliation planes becomes narrower and the subhorizontal stretching lineations become more distinct to the shear zone center. Except for ultramylonites from the center of SZ2 the sheared rocks are medium- to fine-grained S-C mylonites (Lister and Snoke, 1984) showing a sinistral sense of shear. The granodiorites and the mylonites are composed of plagioclase (An _ so), K-feldspar and quartz. From the composition of recrystallized feldspar in the mylonites (albite in plagioclase vs. albite in alkalifeldspar geothermometer of Stormer, 1975) metamorphic temperatures can be roughly estimated as 400-500°C. Mafic minerals in the granodiorite and the mylonites are mainly magnesio-hornblende and biotite. In the granodiorites hornblende occurs as euhedral crystals, sometimes with cores of clinopyroxene. In the mylonites hornblende forms patchy porphyroclasts surrounded by rims of very finegrained biotite and actinolite often developed as trails extending into C-planes. Very small biotite flakes also grew along cracks in the interior of hornblende porphyroclasts (Fig. 5a). Most biotite occurs as rounded, partly recrystallized aggregates (former igneous crystals), a minor amount grew along
Fig. 5. Microphotographs of mylonites of SZ2. (a) Hornblende porphyrcclast (H) in center of SZ2 (sample 88-lg). Biotite (b) growth occurred on the rim, on cleavage planes and in a shear plane; on the rim of the hornblende porphyroclast also some recrystallization of hornblende (h) took place. The shear plane on the upper left side is rather sharp, according to Shimamoto ( 1989) indicating semiductile deformation related to some extent to seismic slip events. (b) Two contrasting sites of biotite recrystallization in a granodioritic mylonite &rnple 88-le). B = large, former igneous grains; (I = small flakes in shear plane. The plagioclase porphyroclasts (at the upper margin of the photograph) shows dynamic recrystallization on the rim along the shear plane. (c) Diopside porhyroclasts (D) in a very fine grained dynamically recrystallized matrix of mainly plagioclase (P), some coarser grained quartz is also present in the matrix, e.g., on the right side above the scale bar.
shear planes (Fig. 5b). The major part of the matrix of the mylonites consists of quartz. The ultramylonites (sample 85-54b) form bands, up to 50 m wide, which run through the “normal” mylonites parallel to the foliation. In these rocks no S-C fabric but extensional crenulation cleavages are developed. The ultramylonites consist mainly of intermediate plagioclase, quartz, diopside, epidote and sphene (Fig. 5~). Plagioclase and, to a minor extent,
E. Scheuber et al. / Tectonophysics 250 11995) 61-87
quartz form the ductile matrix; diopside, epidote and sphene are porphyroclasts. Compared with the “normal” mylonites in the ultramylonites a complete loss of K-feldspar, biotite, hornblende and Fe-ore minerals has taken place which, together with the very fine-grained matrix of the ultramylonites with an average matrix-grain size of N 4 pm, indicates that diffusional mass transfer played an important role in the formation of this rock. In contrast to SZl the incompetent minerals of SZ2 mylonites (quartz and biotite) were much more strongly affected by dynamic recrystallization than the competent minerals (feldspar and hornblende; Table 1; Fig. 4). This ductility contrast indicates that these rocks deformed in a semiductile manner as suggested by Shimamoto (1989) for S-C mylonites. However, the competent minerals are not completely free of dynamic recrystallization, as the amount of dynamic recrystallization of these minerals increases from 5% (88-1~) to 9% (SS-lg) from the margin into the shear zone center. In the ultramylonitic sample (8554b) all plagioclase has recrystallized dynamically. In the S-C-mylonites quartz is the weakest phase, deformed mainly by dislocation creep (indicated by lattice preferred orientations, Scheuber and Andriessen, 1990); quartz is weaker than biotite which forms elongate aggregates (fishes). Again the ultramylonites are an exception in that quartz forms elongate lenses in a plagioclase matrix indicating that plagioclase is weaker than quartz; this is consistent with the findings of experiments reported by Dell’Angelo and Tullis (1982) and Tullis (1990). They described a very distinct strain weakening of feldspar due to dynamic recrystallization by straininduced grain boundary migration. The SZl mylonites contain 19-31% of quartz (exception: ultramylonites 13%). The size of recrystallized new quartz grains is N 45 pm; from this size a flow stress in the order of 40-60 MPa can be calculated (Scheuber, 1994) using the paleopiezometers of Twiss (1977). The quartz content of the mylonites is in the range where a transition takes place from a framework-supported rheology to a matrix-supported rheology (e.g., Jordan, 1988; Handy, 1990). This is consistent with the microstructures of the mylonites: In those samples with a high quartz content (e.g., sample SS-lg: 3 1% qtz) the quartz lenses are more interconnected in matrix form
69
than in the quartz-poor rocks. For example, in sample 88-le (19% quartz) quartz forms elongated lenses oriented parallel to the foliation; however, these lenses are often isolated and do not form an interconnected matrix. Thus, the quartz-poorer samples were stronger during deformation and, consequently, accommodated a smaller fraction of the bulk strain rate of the shear zone than those samples richer in quartz. The finite strain (X/Z ellipticity R) increases from R - 2 (sample 88-1~) to R - 4 (SS-lg) measured using the normalized Fry method (Table 1; Erslev, 1988). The highest strain was accommodated by the ultramylonites with R B- 50 (R,/cp’ method in quartz lenses). However, as most measured particles are disrupted it is not possible to determine which fragments are part of one pre-deformational clast and thus the measured R, value is generally too low. The high strain in the ultramylonites is shown by apatite grains with R, > 1000, in X§ions the length of these grains exceeds 1 cm, whereas their width is only N 10 pm. There are no indications of annealing of the mylonitic fabrics in SZ2. This is inferred from the observation that the grain size of dynamically recrystallized new grains is smallest at sites of stress concentration, e.g., around porphyroclasts or in Cplanes. The varying sizes of new grains in the mylonites of SZ2 indicates that due to a rapid stress drop the peak stress event has “frozen in” the size of the new grains (Knipe, 1989).
4. Results of age determinations Rb-Sr and “OAr/ 39Ar data are listed in Tables 2-4. In Table 5 we summarize the Rb-Sr ages and the “OAr/ 39Ar total degassing ages as well as the plateau ages in the order of decreasing distances from the shear zone centers. In the column “Geological mean age” the mean values of geologically significant @Ar/ 39Ar and Rb-Sr ages are shown. Rb-Sr isochrons (Figs. 6 and 7) were calculated only if the data of at least three mineral components of one hand specimen are available (Table 4). @Ar/ 39Ar total degassing ages were deduced from summing up the data of the incremental heating steps (Table 4). Plateau ages (McDougall and Harrison, 1988) are considered when three or more successive incremen-
70
E. Scheuber et al./ Tectonophysics
tal degassing steps give, within the 2u error limits, the same age value and when the trapped “OAr/ 36Ar ratios are identical to the atmospheric value (295.5). Table Results
250 (1995) 61-87
The trapped 4oAr/ 36Ar ratio is given as the intercept of a line defined by these plateau age steps in a discrimination diagram of 40Ar/ ‘(jAr vs. 39Ar/ 36Ar.
2 of the Rb-Sr
analyses
Sample SZl 87-150
87-153
for whole rock and minerals
material
88-26
88-1 a
88-lc
88-le
88-19
85.54b
Sr [nom1
87Rb/86Sr
20
87Sr/ssSr
16.4 14.1 6.57
471.4 705.3 88.4
0.1008 0.0580 0.2771
0.703496 0.703402 0.704172
* * *
120 45 30
0.25-0.5 0.25-0.5 0.25-0.5
15.3 4.35 5.48 349.8
461.6 786.6 65.5 37.3
0.0961 0.0160 0.2423 27.32
0.703354 0.703258 0.704019 0.761504
f * f +
30 30 180 35
Pig hb bi
0.25-0.4 0.12-0.25 0.25-0.4
20.6 2.51 6.43 307.4
396.7 569.3 31.6 11.6
0.1503 0.0127 0.5902 78.01
0.703795 0.703519 0.704724 0.863298
+ f + r
35 40 35 50
WR LM bi
0.13-0.4 0.13-0.4
55.5 24.6 303.8
360.5 341.7 12.1
0.4455 0.2083 73.33
0.704265 0.703729 0.837958
+ * r
50 50 120
WR LM hb bl
0.25-0.7 0.25-0.7 0.5-0.7
48.4 27.1 3.18 327.2
407.4 416.1 25.9 7.6
0.3424 0.1884 0.3571 126.8
0.703934 0.703681 0.703932 0.936388
+ + * *
80 50 50 120
WR LM bi
0.25-0.5 0.25-0.5
86.6 53.0 556.8
389.3 381.8 20.8
0.6420 0.4004 78.55
0.704652 0.704176 0.847446
+ f -f
70 35 85
WR LM hb bl
0.12-0.25 0.181-0.25 0.12-0.25
78.8 45.6 31.6 552.2
405.8 437.8 27.4 20.9
0.5615 0.3018 3.338 77.50
0.704484 0.703971 0.709735 0.841503
k * * k
150 30 45 50
WR LM bl
0.12-0.25 0.12-0.25
59.7 20.6 439.2
420.3 437 2 24.6
0.4104 0.1346 52.06
0.704236 0.703692 0.799226
+ * +
140 45 70
73.8 61.5 148.0 2.85
403.7 429.2 68.3 391.4
0.5295 0.4137 6.267 0.0210
0.704477 0.704234 0.714746 0.703551
* + + t
50 60 25 70
WR hb bi
sz2 85-58
Rb lppml
0.25-0.5 0.25-0.5
WR LM hb
PM
87-151
grain size (mm)
WR
WR LM hb( + bil WR
0.12-0.25 0.12-0.25
WR = whole rock; plg = plagioclase; LM = light minerals (plagioclase, quartz, K-feldspar); hb = hornblende; bi = biotitc; the grain size (mm) of the analyzed minerals is also given; errors refer to the last digits. Analytical methods: Aliquots (0.1 g) of whole rocks (3-5 kg), plagioclase or light minerals, hornblende and biotite were dissolved in a mixture of hydrofluoric and nitric acid under a pure nitrogen atmosphere. Sr and Rb concentrations were determined using a highly enriched 84Q-spike,and a *‘Rb-spike solution adding to aliquots of the sample solution. Rb and Sr were separated using ion exchange in resin columns. Approximately 0.1 g samples were processed, and the blanks never exceeded 0.50 ng Sr and 0.15 ng Rb. The precision on Sr and Rb concentrations is estimated as 0.1 and OS%, respectively (lo error). The isotope compositions of the samples were measured on a MAT 261 (Finnigan) thermionic mass spectrometer operating in the static mode. The relative transmission of the ion beam between the different faraday cup positions was in the order of 3 X IO-‘. Routinely, the measured NBS 987 standard gave s7Sr/ s6Sr = 0.71022 + 4 (2 u~,,,~) normalized to a value of s6Sr/ ‘*Sr = 0.1194.
E. Scheuber et al./ Tectonophysics
71
250 (1995) 61-87
Sr (e.g., van Blanckenburg et al., 1989). Any deviation from this behaviour is the result of a geological or an artificial disturbance. Such disturbances can be recognized in a staircase age spectrum with the “OAr/ 3gAr method. In the case of biotite there are at least four possibilities of explaining such staircase age spectra: (1) Thermal overprinting which leads to a partial Ar loss (Harrison, 1983); when the thermal pulse is moderate the Rb-Sr biotite system is not affected and as a consequence the 4oAr/ 3gAr total degassing age is lower than the Rb-Sr age. (2) Radiogenic Ar loss during very slow cooling (< l”C/Ma) which also generates staircase age patterns (McDougall and Harrison, 1988). In this case the closure of a system is a temperature range where at high temperatures only a small amount of radiogenie Ar is stored within the mineral and by far the largest amount of Ar is lost by diffusion, at the lower end of this range nearly all radiogenic Ar is kept in the mineral. Chemical bonding of Sr is stronger and therefore Rb-Sr ages are higher than the K-Ar c40Ar/ 39Ar total degassing) ages. (3) Vacuum degassing effects are thought to change the “OAr/ 39Ar distribution. Recent results
Two types of age spectra can be distinguished (Table 5; Figs. 8 and 91: flat plateau ages across a wide range (> 70%) of 3gAr released (e.g., sample 87-150 hornblende, Fig. 8), and staircase patterns sometimes with a plateau-like part of 20-55% of 3gAr released. In relation to the Rb-Sr ages two kinds of plateaus can be distinguished: in SZl samples the plateau age values agree with the Rb-Sr ages and the “OAr/ 3gAr total degassing ages are younger than the Rb-Sr ages (e.g., sample 87-153 biotite, Fig. 8); by contrast, the total degassing ages of SZ2 samples are equal to the Rb-Sr ages whereas the plateau-like parts reveal higher age values (e.g., sample 88-la biotite, Fig. 9). Thus, we have to interpret these staircase age patterns. In thermally undisturbed samples K-Ar biotite ages which are identical to 40Ar/ 3gAr total degassing ages agree with the Rb-Sr ages because both systems shared the same closure temperature of about 300°C (Purdy and Jager, 1976; van Blanckenburg et al., 1989; Hurford et al., 1989). Recently the closure temperature of hornblende K-Ar was revised by Dallmeyer (1978) as 480°C and by Harrison (1981) as 53Or$,“C. Therefore, in a steadily cooling environment we can observe the following age sequence: hornblende K-Ar > biotite K-Ar, Rb-
Table 3 Rb-Sr results
Sample
lsochron I Reference line
SZI 87-153
MSWD
Age
_
e’SrleeSri
20
0.703149
f
25
0.22 0.04
0.703486 0.703490
f +
32 30
i 0.4
2.03
0.703397
f
70
129.3
i
0.3
0.38
0.703326
f
45
WA-LM-bi
129.0
i
0.3
0.65
0.703448
+
60
88-1~
WR-LM-hb WR-LM-bi
133.6 125.3
f 1.2 f 0.3
0.07 0.21
0.703390 0.703434
f f
20 20
88-le
WR-LM-bi
129.4
f 0.3
0.24
0.703448
f
45
88-lg 88-l g + 85-54b
WR-LM-(hb + bi) WR-LM-(hb + bi)WR
126.1 126.1
+ 0.6 f 0.6
0.79 0.39
0.703515 0.703515
k *
40 36
WR-bi
150.3
f 0.3
WR-pig-hb WR-pig-bi
140.5 143.9
f 7.9 f 0.3
85-58
WR-LM-bi
129.1
88-26
WR-LM- bi
B&la
87-151
522
WR = whole rock; pig = plagioclase;
LM = light minerals;
hb = hornblende;
bi = biotite. 2~ error values refer to the last digits.
f f zt
4869 8413 2375
f
4691
i
f
4000
2310 4516
1200 4000
*
*
*
+
f
*
f
4000
1765
1161
750 800
7194
2110
4800 3000
1785
2100
434
*
f f f +
3697
2590
4170
12270
3145
5675
5745
10050
3960
3365
8640
1905
2891
975
990
1005
1020
1040
1070
1120
1130
1180
1240
1320
Total
+
r
f
f
*
*
*
*
+
*
1283 2059
6561
11000 1500
2388
1500 4000
2922
7426
4284
4237
2500
10000
5000
4000
2316
3180
3200 9287
1917
1100
1400
2576
3500
10000
3987
5000
29.7 305
15 100
+
i
+
f
f
f
f
f
f
f
f
*
*
f
*
*
.
0.1540
1100
2700
8600
1000
0.1192
0.1162
0.1140
0.1159
0.1131
1800
0.1285 0.1114
0.1175
3100 3600 7000
0.1115
0.1050
0.1003
0.1118
1100
9000
2400
810
0.3752
1.165
2300
1.323
0.4655
45 3000
1.1
Le 0.26-0.5
950 Sli
0.1248
130
grain
0.1721
0.1504
0.1353
0.1319
0.1292
0.1421
2800
3000
3000
850
550
500
0.2650 0.1612
0.2802
1400 5600
0.6460
3000 4000
1.082
35
*
f
It
*
f
k
2
*
*
*
i
+
k
*
at
- 00002 0.0002
313.2
2.0
24.0
0.0003
15.66
9.17
14.67
145.9
20.2
14.19
12.13
12.78
5.12
3.40
3.43
0.81
88 * 10
415.8
56.2
29.8
I [
174.3
58.3
62.4
24.0
15.93
4.12
3.91
1.31
3.36
4.63
7.28
I.5 mm,
Chile)
zt
i
*
*
f
*
+
f
f
*
*
*
*
*
*
f
6.9
1.1
1.2
1.3
0.90
0.70
0.80
6.0
1.0
0.85
0.80
0.80
0.60
0.60
0.60
0.60
6.5
f
r
5.0 2.5
*
1.9
2.5
*
f
1.0
0.75
* i
0.50
0.50
0.45
0.45
0.45
0.50
it
*
*
*
f f
J = 0.06888
(norhem
0.0003
0.0003
0.0003
0.0006
0.0003
0.0004
0.0003
0.0003
0.0003
0.0010
0.003
0.004
0.015
0.0004
*
*
0.0002
*
mm, +J=OO
0.0006
0.0007
0.0008
0.0004
0.0005
0.0007
0.0006
0.0007
0.0005
0.003 0.0016
*
*
i
*
*
*
*
+
*
f
i
grain size 0.:
Fault Zone, Coastal Cordillera
Ibbro, 30 m N of she ar zone center, 3100
!
4100
70
10 m N of sh ar zone center,
*
nblende,
B7-153
f
i
5510
6589
1180
i
f
820 910
6274
1085
632
5939
1075
315
3186
1070
f
620
2585
1055
f
f
f
f
f
f
undeformed
730
1748
1040
h
2519
10430
1025
2646
965
1000
4253
910
nblende,
789
h
820
87-150
SZI
(*c)
Temp.
Table 4 Argon isotope data for samples from shear zones of the Atacama
10
1.261
1.253
1.272
1.285
1.253
1.309
1.272
1.270
1.229
1.289
1.218
1.196
1.32
1.31
1.10
0.64
1.271
1.267
1.285
1.274
1.250
1.252
1.297
1.25
1.41
1.05
1.32
1.28
1.138
t
0.075
0.070
f
0.027
0.065
f f
0.060
0.070
0.090
0.070
0.060
0.065
0.075
0.080
0.075
0.15
0.20
0.19
0.50
0.020
0.060
0.035
0.040
0.045
0.050
0.060
0.15
0.17
0.40
0.18
0.13
f
_+
*
+
+
*
i
f
*
*
*
*
f
+
f
*
i
f
*
+
It
f
*
*
*
+
40 Ar,d/39Ar,:
17
* f
i *
150
f
*
*
*
f
f
*
*
r
*
f
*
f
f
f
f
*
*
80 3.0
7.5
6.5
8.0
11
8.0
6.0
7.5
9.0
9.5
8.5
17
25
20
60
2.5
7.0
4.0
5.0
5.0
6.0
7.0
20 f
45
20 f
2
151.3
9.0 15
*
f
It
f
152.7
154.1
150.5
156.9
152.6
152.4
147.7
154.7
146.5
143.9
158
157
133
79
152.6
152.1
154.2
152.9
150.1
150.3
155.5
150
168
127
158
153
137.2
Ago
1.6
a.7
4.4
2.6
100.0
99.4
91.8
62.4
77.4
74.6
70.0
23.7
f7.1
12.7
0.5
100.0
86.0
45.0
31.0
15.7
9.8
6.1
5.1
4.2
3.8
3.1
2.0
(%t
C"ill.
j=Ar
21090
47820
25810
22880
22400
21660
5989
11510
855
890
915
940
970
1000
1045
9491
3329
935
970
* f ;
7708
12620
1067
4270
1185
1235
1320
Total
855
590
477.5
299.5
413.5
1055
1110
1190
Total
426.8
5090
4035
815
990
4309
5460
755
775
*
;
*
*
*
*
f
f
*
* f
419.6
1925
715
665
1900
2800
9700
6000
7000
8700
31000
6600
9000
26000
17000
6600
160
-I-
1821
917
3526 3429
3548
4007
7355
3300
4832
6699
8338
4494
*
2.5 40
2.5
0.7
1.3
5
1.5
800
850
950
2100
450
7000
5000
10000
3000
6000
2500
150
1800
6000
2500
1600
4500
2000
f
20.0
12.4
30.0
48.8
21.8
624
779
*
f
*
f
f
*
*
f *
837
f
653
272
26.9
n
0.40
0.05
0.11
0.20
0.09
120
18
140
60
10
0.14
grain size 0.25-0.4
3150
2800
*
f
*
*
*
+
f
*
f
*
*
f
f
80 140
f
*
70
300
450
350 1500
1100
1400
5000
1000
1400
4000
2500
1100
70
2.4
*
7030
9400
5810
7790
9370
6150
3490
9190
2570
7280
2960
1620
5080
300 4000
240
f
*
f i
+
*
*
*
+
f
f
*
+
*
f
*
*
f
*
*
+
i
+
*
*
*
*
f
i
f
*
*
*
+
f
*
*
*
f
*
f
f
f
*
39.8
3.11
58.4
84.3
46.5
100
11s
89
377
77
33.8
f
*
i
i
+
?
*
*
*
*
*
,, J = 0.01395
0.1232
0.1075
0.1081
0.111s
0.1107
0.1626
0.1196
0.1060
0.1137
0.2027
0.2402
0.7789
1.344
4.308 5.063
2.141
1.274
!5 mm,
18.15
2.393
12.23
5.497
11.23
23.57
60.9
28
32.5
37.9
48.3
62.9
46.4
i
16
3.0
19
0.0003
4.0
0.25
3.0
9.0
6.5
60
50
35
210
f
0.0003
0.0003
0.0003
0.0003
0.0004
o.ooo5
0.0002
0.0003
0.0009
0.0016
0.0020
0.004
0.021
0.031
0.007 0.012
f
0.65
0.035
0.35
0.035
0.14
0.90
3.0
12
1.5
2.0
2.5
4.0
2.5
0.090
J = 0.01395
J = 0.06888
mm,
4.326
grain size 0.25-0.5
gain site 0.13-t
75.1 857
700
10000
5500
10000
10000
6000
2550
1000
2500
6500
2000
1500
5000
4000
300
tite. shear zone center
f
f
zt
12350
22230
*
+
*
*
*
*
*
f
*
f
f
1120
87-151
16 440
shear zone CI lter.
*
*
*
*
Y!Z
*
*
*
f
zt
1080
8116
4040
880
1050
1978
835
1498
5558
750
4669
4770
730
1000
561
710
1035
366
660
#7-151
II
*
30720
815
nblende,
42140
775
h
52190
755
TOld
*
27880
715
*
*
459
5254
10 m N of shear zone center,
665
biotite,
585
87-153
1
l-
4409
a4
1189
1352
439.7
130.8
185.8
197.2
308.5
306.3
215.35
314.3
3.2
6.43
31.2
31.8
23.10
106.3
52.1
37.5
6.63
6.07
2.65
2.33
1.53 2.40
0.83
0.29
$689
58.8
231.9
410.9
371.0
368.4
352.0
326.1
510.4
330.5
787.8
579.7
329.5
32.3
k 19
*
*
+
+
f
t
f
*
*
*
f
*
*
+
*
*
f
*
*
*
*
* *
*
f *
*
f
*
f
*
f
*
f
*
It
f
?
*
*
*
2
2.0
2.5
3.0
3.0
3.0
3.0
2.5
3.5
2.5
4.5
3.5
3.0
2.5
30
15
9
9
4.5
3.0
3.0
3.0
3.0
3.0
3.5
6.0
1.1
0.80
1.4
1.5
0.90
4.0
2.5
1.6
0.60
0.60 0.60
0.60
0.60 0.60
0.60
0.60
25
- _-___
* *
1.27 1.26 1.308
f
* *
1.31 1.10
*
6.008
5.898
3.24
6.060
f
i
*
*
f 6.138
f 5.99
f 6.155
6.171
*
*
5.994 6.148
f
4.613
*
f
1.275
1.260
f
*
f
A
1.267
1.286
1.272
1.254
*
*
1.04
1.18
k *
f
1.12 1.04
0.88
*
*
6.160
0.29
f
f
f
f
*
f
+
i
*
*
*
6.21
6.232
6.268
6.361
6.367
6.462
6.320
6.296
6.245
6.223
6.137
f f.
2.18 5.785
0.032
0.60
0.035
0.030
0.055
0.13
0.090
0.085
0.060
0.060
0.070
0.025
0.40
0.17
0.060
0.060
0.050
0.050
0.060
0.055
0.11
0.12
0.30
0.25
0.40 0.25
0.60
0.60
0.020
0.020
0.060
0.035
0.040
0.040
0.055
0.045
0.030
0.045
0.025
0.030
0.045
0.16
*
147.5
138.5
77
142.2
143.9
141.0
140.6
144.3
144.7
144.2
140.7
109.2
151.3
132
157.2
153.0
152.0
154.2
152.7
150.6
156.8
142
152
152
126
125
107 135
36
144.4
145.6
146.0
146.9
148.9
149.1
151.2
f
*
f
f
i
f
f
*
*
*
*
*
f
f
i
f
f
f
*
f
k
*
f
f
f
* *
f
*
f
i
*
*
*
+
i
*
146.3 148.0
f
145.8
*
143.9
t f
52.5 136.0
1.5
14
1.5
1.5
1.8
3.0
2.5
2.5
1.9
1.8
1.9
3.0
45
19
6.5
6.5
6.0
6.0
6.5
6.5
12
14
30
30
30
70 50
30
1.4
5.0
1.9
1.6
1.7
1.7
1.9
1.7
1.5
1.7
1.5
1.5
1.6
3.5
I
6.2
100.0
96.5
70.3
40.6
31.0
28.1
24.1
19.6
13.1
* f f f +
19200
25270
74070
21290
7382
1090
1125
1175
1235
1295
Total
f f
1266 841
Total
)
12
400
350 170
450
400
210
45
160
1
*
1195
550
130
6500
3300
800
f c + * + f f f zt _+ * f +
776
926
2460
7890
11480
11530
25470
19010
24980
26470
73900
13650
11760
730
785
845
905
950
995
1033
1090
1115
1160
1205
1280
Total
f
355
1300
8000
4000
2500
3000
2500
9000
1400
1600
600
120
70
18
30
13.19
4.30
22.32
14.29
10.6
14.72
16.73
66.2
13.37
13.49
16.48
6.85
2.692
2.390
f
f
f
i
*
f
f
i
*
i
f
50.3
808
875 459 _+
* f
+ f
937
*
1094
k
268
f
*
527
274
1.5
300
250 90
300
300
140
18
65
2065
2609
13070
4665
4342
3223
4349
2002
2005
408 1383
230
256
39.2
f
_+
i
+
f
t
f
f
*
zt i
f
f
*
23
15000
7000
400
550
450
1600
250
300
20 110
16
55
2.5
1
I
1 f
f
f
*
it
zt
k
f
ct
0.0003
0.0002
0.0001
0.0001
0.0002
0.0008
0.0026
0.013
/
I
[
200
4.08
5.19
178 4.46
1.62
ze 0.25-0.5
13992
52.0
850.6
760.8
427.7
308.4
200.9
261.1
234.3
233.8
218.8
87.6
24.1
2.1
17.85
1.622
17.59
30.3
12.73
21.44
41.4
17.55
22.15
24.2
11.97
6.296
3.053
1.812
f
i
f
+
*
+
f
i
f
f
f
0.35
0.045
0.12
1.3
0.30
0.40
1.2
0.50
0.90
1.0
0.20
0.090
0.070
f f
0.040
k
I
f
259.4
3992
9.9
451.9
1157.0
620.5
440.4
220.7
311.8
?
*
*
*
f
i
+
k
*
+
*
154.5
*
+ f 324.0
0.40
11
0.55
0.50
11 0.45
22
1.8
3.0
6.0
3.5
2.5
1.8
2.0
1.9
2.0
1.6
1.4
1.3
1.4
J = 0.0139
i
f
i
f f
31.7
7.0
3.2
20
1.9
5.0
4.5
2.5
2.0
1.7
1.9
1.8
1.9
1.8
1.5
1.5
1.4
I 5.569
5.74
5.710
5.745
5.859
5.984
5.762
5.856
5.566
5.514
5.372
4.603
1.516
0.70
f 19 f.
f
f
*
f
f
zt
f
f
f
i
+
f
f
19
5.551
5.12
5.831
5.610
5.684
5.806
5.787
5.612
5.579
5.492
5.306
2.74
1.88
1.51
k
1.12
1.13
1.152
0.83
1.08
1.13
0.42
*
*
+
i
k
i
r
f
*
*
zt
f
f i
*
f
f
i
f
f
f
mm, J = 0.1 6888 f 10
f
i
f
*
*
f
f
zt
+
+
+
k
f
f
.4 mm, J = 0.01395
zone center, grain size 0. 5-t I.7 mm,
0.08908
0.08184 0.08243
0.09295
0.08515
0.3253
0.8564
1.467
0.4623
0.25
0.75
0.25
0.30
1.6
0.35
0.25
2.5
0.50
0.50
0.55
0.12
0.019
* f
0.055
f
)ne center, grain size 0.1:
wdiorite, 5.2 km E ofs ihear zone center,
*
3656
39000 950
*
+
9452
4346
300
200
500
110
130
45
16
8.0
1.1
e, undeformed granor rite, 5.2 km E of she
670
$8-26 bit
f
*
1351
1075
1080
1215
LIC
1556
1045
(
*
1130 1230
f
586
734
1020
T
* f
317
604
910
985
f
1457 +
f
2694
_+
f
1154
2088
i
1198
3227
f
896
+
;
185.3 280
r
29.6
26000
3500
3500
1700
1200
3000
600
700
250
90
30
15
18-26 homlblende, undeformed Q
f
1005
f
+
8719
940 f
*
6904
910
6690
f
4036
875
16080
f
1586
820
12790
*
577
770
1045
f
te, undeformed granod wite. 9 km E of shear
316
bit
710
85-58
sz2
Table 4 (continued)
0.030
0.90
0.035
0.030
0.035
0.035
0.045
0.035
0.040
0.35
0.055
0.12
0.35
0.65
0.11
0.12
0.070
0.20
0.10
0.35
0.25
0.029
0.20
0.035
0.030
0.040
0.040
0.050
0.045
0.045
0.045
0.045
0.080
0.090
0.50
I
I
17
130.6
121
132.5
132.0
133.7
136.4
136.0
132.0
131.3
129.3
125.1
65.7
45.3
36.4
134
136
138.8
101
131
136
52
131.1
135.0
134.3
135.1
140.4 137.6
135.4
137.6
131.0
129.8
128.6
109.0
36.7
i
f
i
f
*
f
k
f
*
f
f
f
f
f
f
f
:
*
f
f
i
f
*
*
* zt
*
f
i
+
f
*
i
f
1.3
2.0
1.4
1.3
1.4
1.4
1.6
1.4
1.4
1.4
1.6
3.0
8.0
15
13
14
8.0
24
12
45
30
1.3
5.0
1.4
1.4
1.5 1.4
1.6
1.5
1.5
1.5
1.5
2.0
2.0
11
1.6 5.2
4.0
100.0
98.0
95.4
93.1
1.0
loo.0
98.6
76.0
55.6
44.7
36.9
31.6
24.8
18.4
12.0
5.6
2.9
0.5
6371 364
4120
1320
1380
Total
I
20460
1120
f
f *
k
f
zt
f
+
*
f
*
f
*
+ k +
6269
8493
15700
28510
19650
820 860
680
890 900
* *
5895
13880
17340
6161
1080
1160
1320
7411
19390
25280
21500
15990
14580
52940
61620
16460
342
940
990
1100
1110
1130
1155
940
975
1005
1055
1280
8441
846
2588
910
528
530
770
Total
zt
+
zt
f
f
f
+_
*
f
f
*
c
=t
fr
*
k
te, mylonite,
840
38-1 e bit
zt
6289 zt
+
9328
*
_+
930 1010
i
f
5443
770
800
* +
f
505
1361 2840
670
730
38-1 c bitHit 8. mylonite.
14200
6355
940
1050
8788
910
6637
54200
890
9365
10740
880
980
2603
820
1020
374
665
850
45
8500
56000
16000
450
10000
9000
5600
2900
700
180
40
25
250
0.9 km
500
4000
1000
280
740
3000
6000
13000
2900
950
2000
1000
160 450
13
1.2 km
400
50
800
4005
4500
1500
1300
1100
1300
40000
630
50
12
I
I
38-1 a bitHit 8, very weakaly defol
*
*
t
*
zk
*
+
*
*
*
+
1.9
130
800
750
250
200
190
200
8200
110
9.0
1.0
f
f
*
f
* *
t
*
+
f
*
*
f
ct f
3.0
86
700
180
130 45
550
1000
2500
500
170
400
180
30 80
1487
8.8
2875
12950
9156
2436
2820
3700
4360
3410
1342
494
274
129.5
66
f
*
k
*
f
+_
zt
*
+
*
+
_+
f
+
*
160
0.90
1500
12000
2900
80
2500
1600
950
500
130
35
12
6.0
26
of shear zone center,
1098
3018
2409
1103 1013
1655
3445
5036
2790
1517
1176
988
258 524
132.7
If shear zone center.
1066
3498
2410
1571
1120
1088
1512
9298
1861
449.9
36.3
ed granodiorite,
i
*
*
f
+
*
f
*
*
*
* *
f * +-
0.12-0.25
+
+
i
*
*
*
*
*
3.71
31.59
0.464
10.3
41.7
105.0
57.5
114.6
31.9
37.8
51.0
227
14.91
6.07
3.557
f
*
*
*
*
f
t
*
*
f
f
*
f
_c
*
0.80
0.020
2.5
2.0
3.0
1.9
0.50
1.5
1.9
2.5
25
0.45
0.15
0.040
0.65
mm,
0.50
0.04
0.16
1.9
0.14
0.20
0.45
350
0.80
0.17
0.27 0.13
0.025 0.43 0.20
mm,
0.14
0.11
0.35
0.30
0.18
0.20
0.14
5.0
0.55
0.30
f c
0.016
r
rain size 0.12-0.25
7.73
11.88
7.93
7.4
5.88
7.75
16.94
539
15.82
8.57
10.58 5.55
1.359 18.92 9.79
ain size
9.01
16.31
13.20
10.43
6.94
6.08
11.04
75.9
13.85
8.95
1.012
I E of shear zone center,
J
.I
L
53.8
158.2
357.2
412.3
246.4
114.9
44.9
20.4
1.8
0.01395
4077
2.1
98.4
207.5
533.8
1825
:
%849
583.0
537.1 832.3
227.6
260.9 116.9
352.4 253.6
330.1
118.8 75.7
71.4
47.2 41.6
0.01395
372.2
386.6
193.4
247.1
146.9
150.9
251.1
264.5
966.3
815.6
14.6
f
f
*
f
*
*
f
f
k
f
*
f
f
zt
f
f
f
f
f *
*
* f
* *
f
f *
*
+ +
19
* 19
k
f
*
*
f
f
It
+
i
*
f
; rai in size 0.12-0.25
22
1.9
1.4
1.7
3.0
10
1.3
1.5
2.5
2.5
1.9
1.5
1.4
1.5
1.3
20
4.5
3.5 4.5
2.0
2.0 1.8
2.5 2.0
2.5
1.8 1.7
1.8
2.0 1.8
3.5
3.0
2.0
2.0
2.0
2.0
2.5
2.5
5.5
5.0
1.9
mm, -r
5.530
5.3
5.622
5.681
5.750
5.864
5.57
5.728
5.728
5.599
5.301
4.641
2.005
1.8
3.5
5.343
5.647
5.640
5.526
5.456 5.434
5.617
5.522 5.603
5.404
5.08
5.209
4.85
1.579 4.13
5.519
5.1
5.698
5.764
5.770
5.771
5.662
5.568
5.615
5.797
5.611
+
zt
*
* It
f
* r
f
*
*
*
* f
*
Lk
*
k
-L
*
k
f
*
f
i
*
It
*
*
f
f
*
*
f
f
f
i
*
f
*
f
f
f
0.01395
5.130
2.17
.I = f
19
0.020
5.0
0.080
0.040
0.025
0.025
0.14
0.050
0.030
0.025
0.035
0.060
0.060
1.3
2.5
0.011
0.035
0.014
0.019
0.080 0.045
0.040
0.030 0.040
0.030
0.11
0.075
0.12
0.070 0.18
0.020
3.0
0.050
0.035
0.060
0.045
0.070
0.070
0.045
0.045
0.020
0.020
0.30 52.1
130.2
125
132.3
133.6
135.2
137.7
131.1
134.7
134.7
131.7
125.0
109.9
48.3
84 43.6
125.9
132.7 132.8
130.1
128.0
132.2 129.5
131.8
130.0
119.9 127.3
122.9
114.7
98.1
38.1
129.9
121
134.0
135.5
135.6
135.6
133.2
131.0
132.1
136.2
132.0
121.1
+
k
f
f
k
*
*
f
*
*
*
*
1.
f
f *
zt
k +
*
2
* *
f
*
f t
f
*
f
r
*
f:
f
*
*
f
f:
f
f
k
f
f
1.3
107
2.1
1.5
1.4
1.4
3.5
1.7
1.4
1.4
1.4
1.7
1.6
59 3.1
1.2
1.3 1.5
1.3
1.5
1.5 2.0
1.5
1.4
3.0 1.4
2.0
3.0
4.0
1.7
1.3
76
1.7
1.5
1.8
1.6
2.0
2.0
1.6
1.6
1.3
1.2
7.0
k +
t t
4005
3491
3272
4227
2564
2063
3935
2119
895
910
935
960
1020
1115
1160
1210 400
600
250
190
1100
350
350
450
650
140
9.5
1.9
f r
2604 3429
f * f
1770 960
t
*
1668
3182
1542
f
i
1973
i
2933
f
3641 3262
f
*
f
1327
362.7
21.1
shs r zone center,
75
500
300
500
190
150
900
300
300
350
500
110
6.0
0.12
gain s
1
1 0.3743
0.0007
1
0.0004
0.1921 r
O.Ogll
0.4377
1
0.0007
0.2569
0.0004
f
35.56
0.005
2.393
0.1361
zt
16.15
0.013
i
* 692.5
+ 59.0
f
* 48.1
116.7
226.8
zt
4.264
30.53
*
k
0.012
34.51
39.26
*
4.723
0.10
33.20
*
f
28.89
i
21.58
10 2.19
0.025
f
8.132
10.12
0.035
0.009
0.005
0.10
f
f
mm, J = 0.068
9.91
6.015
2.967
1.018
t: 0.4-0.7
6.0
2.5
1.4
1.5
4.5
0.45
0.45
0.45
0.45
0.45
0.50
0.50
0.50
0.50
f f
1.137 1.090
f
*
1.093 1.109
f
1.144
f
k
1.146
1.133
f
1.149
+
*
1.147
1.146
*
1.115
i
f
0.731
1.143
+
0.377
0.090
0.010
0.045
0.035
0.015
0.025
0.016
0.030
0.018
0.015
0.014
0.017
0.020
0.017
*
37.7
* *
133.9
f
132.0 136.6
t
* 137.8
138.1
*
*
31.5
1 38.5
*
1 37.1
f
f
38.2
1 38.1
*
134.4
i mt
46 89.3
1.3
5.5
4.0
1.8
3.0
1.9
4.0
2.0
1.8
1.7
2.0
2.5
2.0
11
100.0
91.7
64.6
66.3
36.5
31.6
29.3
25.1
20.0
14.4
9.6
5.7
0.9
Analytical methods: Irradiation (hornblendes: 5 days, biotites: 24 hours) of the samples was carried out under Cd shielding at the Kemforschungszentrum Geesthacht, Gas extraction and purification followed the procedure of Hammerschmidt et al. (1992). The reciprocal sensitivity of the mass spectrometer was determined as 1.85 *0.005X lo- “I cm’/mV using an Ar standard of known volume. The extraction blank for “‘Ar was 2.50f0.25X IO-’ cm’ STP between 400” and 1100°C and increased linearly to 1.5X 10e8 cm’ STP between 1100” and 1500°C. For masses 36, 37 and 38 the measured blanks were = 2 X IO- ” cm? STP, while for mass 39 it was = 3X IO-” cm’ STP “Ar equivalents. The errors in Table 3 are at the 95% confidence level and include components from the statistics of measurements, discdmina6on, blank (usually 20% for j9Ar and 30% for other isotopes), interferences, reciprocal sensitivity, weighing errors and relative measurements of the neutron flux.The following interference corrections were used:“hAr/ “Ar(Ca) = (2.7 + 0.2) X 10-J3XAr/ “Ar(Ca) = (6.0 5 2.0) X 1O-““Ar/ “Ar(Ca) = (6.8 f 0.2) X lo- j4”‘Ar/ “AI = (3.0+3.0)X 10-4i8Ar/‘9Ar(K) = (1.4*0.3) X IO- ““Ar/ ‘“Ar(K) = (6.0 7t 2.0)X IO- “‘Ar(K) = neutron-induced only from K1.U.G.S. constants (Steiger and Ilger, 1977)LP-6 biotite standard (Engels and Ingamells, 1971)
It
k
k
I
A
4471
875
f
1775
f
560.7
800
855
f
303.5
lblende, mylonite,
730
Bf3-lg hc
Table 4 (continued)
E. Scheuber et al./Tectonophysics
250 (1995) 6/-87
77
r 0740 3j E 0 720
87-151 WR plg-bi
87-151 WR-pIg-hb
MSWD
= 148.5 f 7.9 Ma
= 143.9 i 0.3 Ma
= 0.703466 * 32 = 0.22
= 0.703490 f 30
. . : . . . : . ,
0.7030.
0.2
0
: .
4
06
0.4
0
40
20
60
60
Fig. 6. Rb-Sr isochron diagrams of minerals and whole rocks from the shear zone SZl; the insert in the upper left corner of each diagram is an enlarged view of the lower left comer. M refer to data included for age calculation, 0 data not included for age calculation. WR = whole rock; pig = plagioclase: LM = light minerals (plagioclase plus some quartz and K-feldspar); hb = hornblende; bi = biotite.
Table 5 Compilation of all age results (for “Ar/ these steps is indicated)
T
Sample
j9Ar plateau ages the number of consecutive
Number
Fiatmu
Total
of steps
“Ar
A-
releesa
heating steps and the amount of ““Ar released during
Rb-Sr (Ma) mean age (-Ma)
I%1
SZl
_..
87-150
hb
152.6
f
2.4
152.9
f
2.4
12
98.0
87-153
hb
151.3
zk 3.2
151.7
i
3.2
15
99.5
87-l 53 bi
144.4
f
1.4
149.3
f
1.4
5
29.2
150.3
87-151
hb
151.3
f
2.9
151.7
i
2.8
16
99.6
148.5
2 7.9
87-151
bi
138.5
f
1.5
142.8
f
1.4
8
90.3
143.9
f
0.3
44.2
129.1
f _--
0.4 0.3
* 0.3
-i_
s22 85-58
bi
131.1
f
1.3
134.8
f
88-26
hb
137.1
f
7.2
138.6
? 75
88-26
bi
130.6
k
1.3
132.7
55
129.3
f
88-1s
bi
129.9
1.3
134.8
33.9
129.0
* 0.3
hb
f ___
1.0
88-1~
f --_
133.6
f
1.2
88-1~
bi
125.9
+
1.2
131.9
*
1.2
48.3
125.3
f
0.3
88-10
bi
130.2
+ 1.3
133.7
f
1.2
88-10
hb
133.9
*
138.0
k 1.7
1.3
1.3
94.8
zk 0.7
19.9
)_
74.9
129.4 126.1
t 0.3 + 0.6’ * hc
ksnde and biotite
78
E. Scheuber et al./ Tectonophysics 250 (1995) 61-87
obtained by laser spot age determinations indicate a homogenization effect because during vacuum degassing by conventional heating the lattice of hydroxyl-bearing minerals like hornblende (Lee et al., 1991) and biotite (Hodges et al., 1994) breaks down to form oxy-hornblende and oxy-biotite and at higher temperatures metal oxides. These reaction products display an @Ar/ 3gAr distribution different from the original minerals. (4) But also “Ar recoil due to the neutron reaction with K may be responsible for staircase age spectra in two ways. Firstly, “Ar is lost during irradiation from the mineral and as a consequence the age seems to be too old (Hess and Lippolt, 1986). Secondly, “Ar recoil from only slightly altered biotite (Hess et al., 1987) to K-poor mineral phases like chlorite or Fe-Ti oxides and titanite which sometimes can only be identified by TEM methods leads to age spectra whose inferred ages are geologically too high and thus not meaningful even if well-defined plateau ages are given. Thus, only independent isotope age determinations corroborate the reliability of the ages. In this respect it is plausible to test our “OAr/ 3gAr results by Rb-Sr determinations whenever it was possible. 4.1. Age results of the amphibolite facies shear zone (SZI J 4.1 .I. Hornblende
From the undeformed gabbro through the margin to the center of the shear zone hornblende yielded remarkably constant plateau ages of 152.9 + 2.4 Ma (sample 87-150) 151.7 f 3.2 Ma (87-153), and 15 1.7 f 2.9 Ma (87- 15 l), respectively (Fig. 8). These ages agree with the total degassing ages (Tables 4 and 5) and, together with the amount of “Ar released (98-99.6%, Table 5), they indicate that no significant Ar-loss has occurred since the closure of the system. Rb-Sr data on hornblende from the shear zone center (87-151) gave an age of 148.5 + 7.9 Ma (Table 3), which is in agreement with the 40Ar/ 39Ar ages. However, from the undeformed and moderately deformed samples (87-150, 87-153) no Rb-Sr homblende age information can be obtained. 4.1.2. Biotite The 4oAr/ 39Ar as well as Rb-Sr ages of biotite display a pattern completely different from the hom-
blendes. Biotite from the weakly deformed sample 10 m north of the shear zone center (87-153) gave a Rb-Sr age of 150.3 _t 0.3 Ma which agrees with the “OAr/ 39Ar plateau age of 149.3 + 1.4 Ma (Fig. 8). But the lower total degassing age of 144.4 + 1.4 Ma together with the staircase pattern of this age spectrum reveal an Ar-loss of 5.3%. From the shear zone center (87-151) still younger age values have been determined on strongly recrystallized biotites which yielded a Rb-Sr age of 143.9 f 0.3 Ma (Fig. 6) and a plateau age of 142.8 + 1.4 Ma (Fig. 8). In this sample also an Ar loss is indicated by a younger total degassing age of 138.5 -+_1.5 Ma. The finding that the Rb-Sr biotite ages are equal to the plateau ages but older than the total degassing ages suggests thermal overprinting which can also be inferred from the statically recrystallized fabrics of the samples (Section 3). Radiogenic Ar is lost more easily than Sr, and this Ar loss can be detected only by the *Ar/ 3gAr method. The staircase age pattern is not a result of very slow cooling (< l”C/Ma), because in a sample, collected close to SZl, fission track age dating on apatite (Scheuber and Andriessen, 1990) indicates an age of 110 Ma. Taking the closure temperature of apatite (100°C) and of biotite (300°C) and their age values into account an integrated cooling rate of S.O”C/Ma is inferred. This value is approximately one order of magnitude higher than that of a very slowly cooling enviroment. Obviously, vacuum degassing and recoil does not affect the 40Ar and 3gAr distribution since the plateau ages are in accordance with the Rb-Sr ages. 4.2. Age results from
the greenschist
facies
shear
zone (SZ2) 4.2.1. Hornblende
Euhedral hornblende from the undeformed granodiorite gave a total degassing age of 137.1 f 7.2 Ma (sample 88-26, Fig. 9). From the age spectrum the highest obtained age value amounts to I38 + 7.5 Ma. Spotty homblendes from the shear zone center confirm this age: the age spectrum of hornblende (88-lg) reveals an age of 138.0 + 1.7 Ma (Fig. 9). But the staircase pattern also indicates an Ar loss of 5.3%. Therefore the total degassing age is lower at 133.9 + 1.3 Ma. The Rb-Sr data of the undeformed rocks (Fig. 7) show a high degree of isotopic disequi-
E. Scheuber et al./ Tectonophysics 250 (1995) 61-87
bi _,,=.-
,
0
80
60
40
20
0
20
40
60
loo
80
120
$10
“Rb?%r
“Rb/%
1
m-.-_r”.“-
0.711
0
20
40
80
60
100
0
1
4
3
2
“Rbl%r
“RbPCN
0.820
0.760
I
t
6 0.780 % 3j 0 0.740
02
04
0.6
_'
= 129.4 L 0.3 M8 = 0.70344a f 4s MSWD 10.24
0.700 0
20
SO
40
100
60
*‘RW”Sr
o
I
2
3
4
0
10
20
30
40
50
8’RbPSr
5
6
7
“Rbl%r Fig. 7. Rb-Sr isochron diagrams
of minerals and whole rocks from the shear zone SZ2; for explanation
see Fig. 6.
60
80
E. Scheuber et al./
Tectonophysics 250 (1995) 61-87
librium between hornblende and the other constituents. In the 87Sr/ 86Sr vs. 87Rb/ 86Sr diagram (Fig. 7) and in the reduced isochron plot (Fig. IO) hornblende shows isotope disequilibrium; hence no Rb-Sr age information could be gained. By contrast, in the mylonite sample (88-1~) whole rock, the light mineral fraction and hornblende define an isochron whose slope corresponds to an age of 133.6 + 1.2 Ma with 87Sr/ “Sr initial ratio of 0.703398 + 20 (MSWD = 0.07). 4.2.2. Biotite The undeformed granodiorite (sam les 85-58 and B 88-26) yielded identical Rb-Sr and Ar/ 39Ar total degassing ages between 13 1.l and 129.1 Ma (Table 5; Figs. 7 and 9). From the weakly deformed gran-
152.9 f 2.4 Ma
odiorite to the shear zone the biotite ages drop from 129.9-129.0 Ma (88-la) to 125.9-125.3 Ma (88-1~). The same age (126.1 + 0.6 Ma, Sri = 0.703515 + 40) results from an isochron defined by whole-rock, light minerals and hornblende containing - 30 wt.% small (< 20 pm> biotite flakes (88-lg). Within the shear zone sample 88-le still preserves the higher age of 130.2 + 1.3 Ma cmAr/ 39Ar> and 129.4 + 0.3 Ma (Rb-Sr), respectively. Comparing the 4oAr/ 39Ar and Rb-Sr age results from SZl and SZ2 we find in SZ2 a sequence of age values completely different from those of SZl. In SZ2 @Ar/ 39Ar plateau ages are generally higher than Rb-Sr ages, whereas the total degassing ages agree perfectly with the Rb-Sr ages (Table 5). According to Roddick et al. (1980) and Foland (1983)
87-150 hb >
total degassing age: 152.6 i 2.4 Ma 20
40
60
80
100
per cent of “Ar released
3 180-z.
67-153 biotiie
87-l 53 hb
P $ H m
140-*-
151.7 f 3.2 Ma 5’
100.
0
>
I’y
:
total &gassing age: 151.3 ) 3.2 Ma 20
40
60
80
100
0
20
per cent of “Ar released
:
;
:
total degassing age: 144.4 f 1.4 Ma
40
80
80
j 100
per cent of “Ar released
7 z
180
0 E E
140 total degassing age: 138.5 f 1.5 Ma
H lQ loo 0
20
40
60
80
100
per cent of “Ar released
Fig. 8. “Ar/
39Ar age spectra of samples from the amphitdite
20
40
60
per cent of “Ar released
facies shear zone (SZl).
80
100
E. Scheuber et al. / Tectonophysics
250 (I 995) 61-87
81
180 c 160 5 140 8l n 120 80-r
I!
60 -I* 40 + 0
total degassing age: 137.1 f 7.5 Ma
E
100
5a 0
8o 60
40
60
80
total degassing age: 133.9 f 1.7 Ma
40 4 0
1 20
;j
100
I 20
per cent of “‘Ar released
40
60
80
100
per cent of 3eAr released
88-26 biotite
132.7 f 0.7Ma
total degassing age: 130.6 f 1.3 Ma
total degassing age: 131.1 t 1.3 Iha 20
40
60
80
20
100
per cent of “Ar released
40
60
80
100
per cent of “Ar released
5
% 80 60--
‘i assing age: 129.9 f 1.3 Ma 40
60
80
n 100
0
per cent of “Ar released
60
total degassing age: 125.9 f 1.2 Ma
40+ 20
40
60
80
100
per cent of “Ar released
total degassing age: 130.2 f 1.3 Ma
20
40
60
80
100
per cent of “Ar released
Fig. 9. “OAr/ j9Ar age spectra of samples from the greenschist
elevated plateau ages were related to excess argon; however, in their cases the total degassing ages of micas were higher than the Rb-Sr ages due to the incorporation of excess argon. By contrast, in our samples the concordance of total degassing and Rb-Sr
facies shear zone (SZ2).
ages excludes both explanations: excess argon or loss of 39Ar during neutron irradiation. Although our biotite samples were microscopically homogeneous the only explanation for the Rb-Sr and “OAr/ 39Ar age discrepancy we can think of is that given by
82
E. Scheuber et al./
Tecionophysics 250 (1995161-87
55
87-151
"0 '- 35 g VI r x 15 .5
‘g
WR
0.01
0.1
bi
$b
il. 0.01
0.1
1
10
100
%b/-Sr
0.1
1
10
1
10
100
10
100
*'Rb/%r
100
1000
0.01
0.1
1 B'Rb/%ir
87Rb/%r
Fig. 10. Deviation from the isochrons (solid lines) of samples from SZl (87-153 moderately deformed, 87-151 strongly deformed) and from SZ2 (88-26 undeformed, 88-lg S-C mylonite and 85-54b ultramylonite). Deviation = [(s’Sr/ *6Sr,)/(87Sr/ s6Sr,, - l}] x lo5 where the suffix m and BF denotes the measured and the best fit values, respectively. WR = whole rock; bi = biotite; hb = hornblende; plag = plagioclase; LM = light minerals.
Hess et al. (1987).These authors have demonstrated that microscopically perfect biotites with only a small K-deficit due to altered inclusion phases like chlorite and Fe-Ti oxide s, only visible with TEM methods, reveal staircase age patterns and plateau ages due to recoil 39Ar which is redistributed from K-rich lattice sites to K-poor ones. The plateau ages are higher than the geological constraint ages and a K-content of 8% and more seems to be necessary to circumvent such problems. As a further consequence of the ‘“Ar redistribution we find sometimes very low age values in the first degassing steps which, however, do not indicate Ar loss by thermal overprinting, because, from the mass balance point of view, it is obvious that ,me plateau-like age values which are higher than the Rb-Sr ages have to be compensated by low apparent ages in the low-temperature degassing steps. From the observation that the K-contents of the biotites from SZ2 were always smaller than 7.75
wt.% (microprobe analyses) and that the Rb-Sr ages was lower than the plateau age we are probably faced with the problem of 39Ar redistribution. Therefore, in the case of biotites from SZ2 we consider total degassing ages only. 4.3. Redistribution
of Sr isotopes
The diagrams of Fig. 10 show the deviation in parts per lo5 of Rb-Sr data from the isochrons. In SZl from the moderately deformed sample (87-153) plagioclase and hornblende show a deviation from the whole-rock-biotite reference line of some 110 and 500 ppm, respectively, hence these data were omitted. By contrast, the data of sample 87-151 (shear zone center) only show deviations of 3-39 ppm indicating a very strong isotopic homogeneity of the rock. In SZ2 a similar homogenization of the Sr isotopes can be observed. In the undeformed sample (88-26) hornblende shows a deviation of
E. Scheuber et al. / Tectonophysics
- 71 ppm from the isochron; by contrast, in the mylonitic sample (88-lg) the deviation of homblende + biotite is reduced to - 1 ppm. Furthermore, from an ultramylonitic band from the shear zone center (sample 8554b) we obtained a very low 87Rb/ 86Sr ratio of 0.02 (Table 1); this agrees with the observed lack of any K-feldspar, biotite and hornblende. The measured 87Sr/ 86Sr ratio amounts to 0.703551 &-70. This sample fits the isochron of sample 88-lg perfectly, indicated by a deviation of only - 2 ppm. In particular a recalculation of the isochron (88-lg hb + bi, LM, WR) including 85-54b whole rock gave an isochron age of 126.1 f 0.6 Ma (MSWD = 0.39) with an initial 87Sr/86Sr ratio of 0.703515 f 18.
5. Discussion: geological interpretation of the ages 5.1. Atnphibolite facies shear zone (SZI 1 The high degree of consistency of the age data reported in the previous chapters is illustrated by the fact that 4oAr/ 39Ar and Rb-Sr methods gave identical ages and that each age range has been obtained from at least two samples; e.g., all biotites of undeformed samples in SZ2 yielded identical 40Ar/ 39Ar and Rb-Sr ages. Uniform *Ar/ 39Ar plateau ages of 153- 152 Ma (Table 5) have been obtained from igneous as well as from completely recrystallized homblendes irrespective of the state of strain across the shear zone. Thus, we interpret the hornblende age as a cooling age due to the regional cooling of the gabbro and its surrounding rocks. Hornblende dates the time when the shear zone together with its host rock passed the isotherm of the hornblende closure temperature; on the other hand, the age of the first deformational event which produced the recrystallized homblendeplagioclase fabric is postdated. Furthermore, the biotites from the moderately deformed sample are about 2-3 Ma younger (149.8 Ma, geological mean age) than hornblende (Table 5); this is in accordance with the concept of closure temperatures where homblende (T, _ 530°C Harrison, 1981) closes prior to biotite (T, - 300°C Purdy and Jager, 1976; Harrison, 1985; van Blanckenburg et al., 1989; Hurford et al., 1989). From the age difference between hornblende
250 (1995) 61-87
83
and biotite the rate of cooling can be calculated as 70- 11 S”C/Ma. In the shear zone center (87-15 1) the Rb-Sr and the @Ar/ 39Ar ages of biotite are completely reset due to the recrystallization of this mineral. The ages amount to 143.9 &-0.3 Ma (Rb-Sr) and 142.8 + 1.4 Ma (“OAr/ 39Ar). This complete age resetting and the homogenization of the Rb-Sr system indicates that the time of deformation 143.4 + 0.3 Ma ago is determined and that it is not a mixed age. From the fact that biotite was reset at 143.4 Ma whereas homblende shows no resetting, it can be concluded that this second deformation was not strong enough or that the temperatures were not sufficient to affect hornblende. The Ar-loss of 3.0% in sample 87-15 1 is interpreted as being the result of static recovery of the microstructure which, however, was not sufficient to disturb the Rb-Sr system. It is important that this recovery affected only the shear zone center and not the moderately deformed sample; thus, static recovery was rather due to a relatively slow stress drop after deformation than to later heating, as in the latter case static recovery should also have affected the moderately deformed sample 87-153. In summary, from the age relations in SZl the following conclusions on the course of crystallization and deformational events can be drawn: (1) Before 152- 150 Ma (a) the intrusion of the gabbroic protolith took place, followed (b) by the development of the recrystallized plagioclase-hornblende fabric due to sinistral strike-slip movements, and (c) by the static growth of biotite (only detectable in the moderately deformed sample (87-153). (2) At N 143 Ma the second deformation (dip-slip normal movements) occurred which led to the recrystallization of biotite in the shear zone center which (3) was also subject to minor static recovery after 143 Ma. 5.2. Greenschist facies shear zone (SZ2.J The oldest age value of SZ2 (139- 138 Ma) is yielded by the @Ar/ 39Ar spectra of hornblende from the undeformed granodiorite as well as from homblende porphyroclasts from the interior of the shear zone. This is interpreted as the cooling age of the igneous rock below the closure temperature of homblende. Rb-Sr hornblende ages show partly negative or geologically unreasonable age values; they do not
84
E. Scheuber et al. / Tectonophysics
give an age information because they are not in isotopic equilibrium with other rock constituents. The biotite ages of - 130 Ma (Rb-Sr and “OAr/ 39Ar) are also interpreted as cooling ages of the undeformed granodiorite. A cooling rate of - 2030”C/Ma can be calculated from the age difference between biotite and hornblende. The younger biotite ages of 126- 125 Ma (Rb-Sr and 4oAr/ 39Ar) are restricted to the mylonites; this suggests that this age is due to resetting during mylonitization. From the fact that resetting was associated with a homogenization of the Sr isotope system it is concluded that resetting was complete in the samples which gave 126- 12.5 Ma. Thus, the reset age is the deformation age and not a mixed value. Although resetting was complete in single samples not all mylonites show the same degree of resetting. Sample 88-le shows the original cooling age of biotite ( - 129.8 Ma); this is interpreted as a relic age (see below). The age relations in SZ2 can be summarized as follows: (1) cooling of the igneous protolith below c of hornblende occurred at - 138 Ma; (2) cooling below c of biotite at - 130 Ma; (3) shear deformation took place at 126- 125 Ma, and no static recovery after deformation can be observed. 5.3. Age resetting by deformation According to the criteria outlined in Section 1 in each shear zone the age of one deformational event could be determined because (1) the same mineral species is younger in the deformed than in the undeformed rocks, and (2) the recrystallized minerals are in isotopic equilibrium. In SZl the first deformation took place above the closure temperature of the dated minerals; this deformational age is therefore obscured by the regional cooling. Consequently only a minimum age for this deformation could be determined. On the other hand, the younger deformation (at - 143 Ma) could be dated as it is restricted to the deformed rocks. Deformation thus started below the closure temperature of the host rock. The homogenization of Sr shows that resetting was complete and thus the obtained age is not a mixing age. Further resetting after deformation due to partial static recovery had only the minor effect of some loss of At-. In SZ2 deformation also started below T,
250 (1995) 61-87
of the protolith and age resetting was also complete, indicated by Sr homogenization in the shear zone center. Furthermore, in this shear zone no postdeformational loss of Ar occurred due to the rapid stress drop after deformation. The patterns of age resetting and homogenization of Sr isotopes can be interpreted in terms of the influence of finite strain, strain rate, stress, deformation mechanisms and temperature: resetting and homogenization is most complete in the most strongly deformed samples from the centers of the shear zones. This suggests an influence of the finite strain on the resetting behaviour. Weakly or moderately deformed samples may show relic ages or only partial resetting. However, not only the finite strain but also strain rate and stress and probably deformation mechanisms have influenced the age resetting. One observation which indicates that stress and strain rate influenced age resetting is the relic age of the mylonitic sample 8%le which has a lower quartz content than the other samples except for the ultramylonite 8%54b (Table 1). The low quartz content resulted in a higher strength during deformation, and, consequently, a lower strain rate. Another observation illustrating the influence of deformation mechanisms is the age resetting and complete Sr-isotope homogenization in the ultramylonite sample 85-54b. The very fine grain size and the compositional changes that affected this rock suggest that diffusional mass transfer was operative as one deformation mechanism, which is a very probable reason for isotopic homogenization. Although age resetting and homogenization can be explained by deformation alone, a further influence of shear heating cannot be excluded. In SZl shear heating can be concluded from the presence of metamorphic orthopyroxene which is limited to the center of SZl. In SZ2 the increasing degree of dynamic recrystallization of feldspar in the more strongly deformed samples may be due to an increase of temperature.
6. Geological implications We have obtained tight age constraints for the timing of movements along the Atacama Fault Zone. In SZl the cooling age of homblendes ( - 152 Ma) -which is very common as an intrusion age in the
E. Scheuber et al. / Tectonophysics 250 (1995) 61-87
Coastal Cordillera (Maksaev et al., 1988; Her& and Marinovic, 1989; Boric et al., 1990)--corresponds to the Araucanian tectonic phase which took place in the Kimmeridgian (Riccardi, 1988), and which resulted in a > 60” angular unconformity between the Jurassic arc lavas and very coarse overlying conglomerates of Kimmeridgian to Early Cretaceous age. On a regional scale the Kimmeridgian deformations in the magmatic arc are contemporaneous with a change from marine to continental sedimentation in the eastward adjacent backarc area @rinz et al., 1994). The age resetting at N 143 Ma has more local importance; it corresponds to vertical movements around the Pluton Cristales. The deformation age of SZ2 gives the age of the strongest sinistral movements along the AFZ. In contrast to the Late Jurassic shear zones, the occurrence of Early Cretaceous mylonites is longitudinally much more constant. The ages of SZ2 coincide with several fission track data obtained from the Coastal Cordillera (130- 100 Ma; Maksaev et al., 1988; Maksaev, 1990; Scheuber and Andriessen, 1990). The deformation age in SZ2 thus coincides with the onset of cooling and uplift of the Jurassic-Early Cretaceous magmatic arc. Uplift and cooling were followed, at about 110 Ma, by the eastward shift of the arc into the Longitudinal Valley (Scheuber et al., 1994). Acknowledgements This work was part of the project “Mobility of Active Continental Margins” and also of the Sonderforschungsbereich 267 (“Deformation Processes in the Andes”) supported by the Deutsche Forschungsgemeinschaft. Thanks to Prof. Reutter, Berlin, and to the colleagues from the Universidad Catolica de1 Norte, Antofagasta for joint field work and discussions. We express gratitude to the staff of the Kemforschungszentrum Geesthacht for facilitating fast neutron activation. We are indebted to Dr. S. Teufel and M. Feth for supporting the experimental work. References Boric, R., Diaz, F. and Maksaev, V., 1990. Geologia y yacimientos metaliferos de la Regi6n de Antofagasta. Serv. Nat. Geol. Min. Bol. 40, 246 pp.
85
Brown, M., Diaz, F. and Grocott, J., 1993. Displacement history of the Atacama fault system 25”OO’S-27%0’S, northern Chile. Geol. Sot. Am. Bull., 105: 1165-I 174. Bucher, K. and Frey, M., 1994. Petrogenesis of Metamorphic Rocks. Springer, New York, NY, 6th ed., 318 pp. Chermiak, D.J. and Watson, E.B., 1992. A study of strontium diffusion in K-feldspar, Na-K feldspar and anorthite using Rutherford Backscattering Spectroscopy. Earth Planet. Sci. Lett., 113: 41 l-426. Dallmeyer, R.D., 1978. 4oAr/ r9Ar incremental release ages of hornblende and biotite across the Georgia Inner Piedmont: Their bearing on late Paleozoic-early Mesozoic tectonotherma1 history. Am. J. Sci., 278: 124-149. Dell’Angelo, L.N. and Tullis, J., 1982. Textural strain softening in experimentally deformed aplite. Trans. Am. Geophys. Union, 63: 438. Derby, B., 1990. Dynamic recrystallization and grain size. In: D.J. Barber and P.G. Meredith (Editors), Deformation Processes in Minerals, Ceramics and Rocks. (Mineralogical Society Series.) Unwin-Hyman, London, pp. 3.54-364. Diaz, M., Cordani, U.G., Kawashita, K., Baeza, L., Venegas, R., He&, F. and Munizaga, F., 1985. Preliminary radiometric ages from the Mejillones Peninsula, Northern Chile. Comunicaciones, 35: pp. 59-67. Dodson, M.H., 1973. Closure temperature in cooling geochrono logical and petrological systems. Contrib. Mineral. Petrol., 40: 259-274. Dodson, M.H., 1976. Kinetic processes and thermal history of slowly cooling solids. Nature (London), 259: 55 l-553. Dodson, M.H., 1979. Theory of cooling ages. In: E. JIger and J.C. Hunziker (Editors), Lectures in Isotope Geology. Springer, Berlin, pp. 196202. Drury, M.R. and Urai, J.L., 1990. Deformation-related recrystallization processes. Tectonophysics, 172: 235-253. Engels, J.C. and Ingamells, CO., 1971. Information sheet 1 and 2, LP-6 Bio 40-60 mesh. U.S. Geol. Surv., Menlo Park. Erslev, E.A., 1988. Normalized center-to-center strain analysis of packed aggregates. J. Struct. Geol., 10: 201-209. Foland, K.A., 1983. 40Ar/39Ar incremental heating plateaus for biotites with excess argon. Isot. Geosci., 1: 3-21. Frost, H.J. and Ashby, M.F., 1982. Deformation-mechanism Maps. Pergamon Press, Oxford, 166 pp. Garcia, F., 1967. Geologia de1 Norte Grande de Chile. Simposium sobre el Geosinclinal Andino. Sot. Geol. Chile, 3: l-138. Getty, S.R. and Gromet, L.P., 1992. Geochronological constraints on ductile deformation, crustal extension, and doming about a basement-cover boundary, New England Appalachians. Am. J. Sci., 292: 359-397. Gonzalez, G., 1990. Patrones estructurales, modelo de ascenso, emplazamiento y deformation de1 Pluton de Cerro C&tales, Cordillera de la Costa al sur de Antofagasta, Chile. Memoria de Titulo, Universidad Catolica de1 Norte, Antofagasta, pp. 135 (unpubl.). Gonzalez, G., 1993. Tectonic interpretation of mesoscopic structures in a high-strain shear zone of the Atacama Fault System, Coastal Range, northern Chile. Abstr. Volume, Second ISAG, Oxford, pp. 183- 185.
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Grocott, J., Brown, M., Dallmeyer, R.D., Taylor, G.K. and Treloar, P.J., 1994. Mechanisms of continental growth in extensional arcs: An example from the Andean plate-boundary zone. Geology, 22: 391-394. Halpem, M., 1978. Geological significance of Rb-Sr isotopic data of northern Chile crystalline rocks of the Andean orogen between 23” and 27’ S. Geol. Sot. Am. Bull., 89: 522-532. Hammerschmidt, K., Diibel, R. and Friedrichsen, H., 1992. Implication of “OAr/ s9Ar dating of early Tertiary volcanic rocks from the North Chilean Precordillera. Tectonophysics, 202: 55-81. Handy, M.R., 1989. Deformation regimes and the rheological evolution of fault zones in the lithosphere: the effects of pressure, temperature, grainsize aund time. Tectonophysics, 163: 119-152. Handy, M.R., 1990. The solid-state flow of polymineralic rocks. I. Geophys. Res., 95: 8647-8661. Harrison, T.M., 1981. Diffusion of J”Ar in hornblende. Contrib. Mineral. Petrol., 78: 324-331. Harrison, T.M., 1983. Some observations on the interpretation of 4oAr/ s9Ar age spectra. Isot. Geosci., 1: 319-338. Harrison, T.M., 1985. Diffusion of “OAr in biotite: temperature, pressure and compositional effects. Geochim. Cosmochim. Acta, 49: 246-2468. Her&, M. and Marinovic, N., 1989. Geocronologia y evolucidn de1 batolito Vicuiia Mackena, Cordillera de1 la Costa, Sur de Antofagasta (24-25”s). Rev. Geol. Chile, 16: 31-49. Hess, J.C. and Lippoh, H.J., 1986. Kinetics of Ar isotopes during neutron irradiation: ‘9Ar loss from minerals as a source of error in 4oAr/ 39Ar dating. Chem. Geol. (Hot. Geosci. Sect.), 59: 223-236. Hess, J.C., Lippolt, H.J. and Wirth, H., 1987. Interpretation of 40Ar/39Ar spectra of biotites: evidence from hydrothermal degassing experiments and TEM studies. Chem. Geol. (Hot. Geosci. Sect.), 66: 137-149. Hodges, K.V., Hames, W.E. and Bowring, S.A., 1994. 40Ar/ 39Ar age gradients in micas from high-temperature-low-pressure metamorphic terrain: evidence for very slow cooling and implication for the interpretation of age spectra. Geology, 22: 55-58. Hubbard, M.S. and Harrison, T.M., 1989. JOAr/‘9Ar age constraints on deformation and metamorphism in the Main Central Thrust zone and Tibetan slab, eastern Nepal Himalaya. Tectonics, 8: 865-880. Hurford, A.J., Flisch, M. and Jagger, E., 1989. Unravelling the thermo-tectonic evolution of the Alps: a contribution from fission track analysis and mica dating. In: M.P. Coward, D. Dietrich and R.G. Park (Editors), Alpine Tectonics. Geol. Sot. Spec. Publ., 45: 369-398. Jaillard, E., Soler, P., Carher, G. and Mourier, T., 1990. Geodynamic evolution of the northern and central Andes during early to middle Mesozoic times: a Tethyan model. J. Geol. Sot. London, 147: 1009-1022. Jordan, P., 1988. The rheology of polymineralic rocks-An approach. Geol. Rundsch., 77: 285-294. Klapcr, E.M., 1990. Reaction-enhanced formation of eclogite-
facies shear zones in granulite-facies anorthosites. In: R.J. Knipe and E.H. Rutter (Editors), Deformation Mechanisms, Rheology and Tectonics. Geol. Sot. Spec. Publ., 54: 167-173. Kligtield, R., Hunziker, J., Dallmeyer, R.D. and Schamel, S., 1986. Dating of deformation phases using K-Ar and 4oAr/ j9Ar techniques: results from the Northern Apennines. J. Shuct. Geol., 8: 781-798. Knipe. R.J., 1989. Deformation mechanisms-recognition from natural tectonites. J. Struct. Geol., 11: 127-146. Lee, J.K.W., Onstott, T.C., Cashman, K.V., Cumbest, R.J. and Johnson D., 1991. Incremental heating of hornblende in vacua: implication for “Ar/ ‘9Ar geochronology and the interpretation of thermal histories. Geology, 19: 872-876. Leloup, P.H., Harrison, T.M., Ryerson, F.J., Wenji, C., Qi, L., Tapponnier, P. and Lacassin, R., 1993. Structural, petrological and thermal evolution of a Tertiary ductile strike-slip shear zone, Diancang Shari,, Yunnan. J. Geophys. Res., 98: 67156743. Listcr, G.S. and Snoke, A.W., 1984. S-C mylonites. J. Struct. Geol., 6: 617-638. Maksaev, V., 1990. Metallogeny, geological evolution, and thermochronology of the Chilean Andes between latitudes 21” and 26” South, and the origin of major porphyry copper deposits. Ph.D. Thesis, Dalhousie Univ., Halifax, N.S., 554 pp. Maksaev, V., Boric, R., Zentilli, M. and Reynolds, P.H., 1988. Metallogenetic implications of K-Ar, ‘k’Ar-39Ar, and fission track dates of mineralized areas in the An&s of northern Chile. V Congr. Geol. Chileno Actas, 1: B65-B86. McDougall, I. and Harrison, T.M., 1988. Geochronology and Thermochronology by the @Ar/ s9Ar method. Oxford Monogr. Geol. Geophys. 9, 212 pp. Poirier, J.-P., 1985. Creep of Crystals. Cambridge Univ. Press. New York, NY, 260 pp. Prinz, P., Wilke, H.G. and Van Hillebmndt, A., 1994. Sediment accumulation and subsidence history in the Mesozoic marginal sea in N-Chile. In: K.-J. Reutter, E. Scheuber and P. Wigger (Editors), Tectonics of the Southern Central Andes. Springer, Berlin, pp. 219-233. Prior, D.J., Knipe, R.J. and Handy, M.R., 1990. Estimates of the rates of microstructural changes in mylonites. In: R.J. Knipe and E.H. Rutter (Editors), Deformation Mechanisms, Rheology and Tectonics. Geol. Sot. Spec. Publ., 54: 309-319. Purdy, J.W. and Jager, E., 1976. K-Ar ages on rock-forming minerals from the Central Alps. Mem. 1st. Geol. Univ. Padova, xxx: l-31. Riccardi, A.C., 1988. The Cretaceous System of Southern South America. Geol. Sot. Am. Mem., 168: 168. Roddick, J.C., Cliff, R.A. and Rex, D.C., 1980. The evolution of *Ar-s9Ar analysis. Earth excess argon in alpine biotites-a Planet. Sci. Lett., 48: 185-208. Scheuber, E., 1994. Tektonische Entwicklung des nordchilenischen aktiven Kontinentalrandes: Der Einflul3 von Plattcnkonvergenz und Rheologie. Geotektonische Forsch., 81: l-131. Scheuber, E. and Andriessen, P.A.M., 1990. The Kinematic and geodynamic significance of the Atacarna Fault Zone, northern Chile. J. Struct. Geol., 12: 243-257.
E. Scheuber et al./Tectonophysics Scheuber, E. and Reutter, K.-J., 1992. Magmatic arc tectonics in the Central Andes between 21” and 25”s. Tectonophysics, 205: 127-140. Scheuber, E., Bogdanic, T., Jensen, A. and Reutter, K.-J., 1994. Tectonic development of the North Chilean Andes in relation to plate convergence and magmatism since the Jurassic. In: K.-J. Reutter, E. Scheuber and P. Wigger (Editors), Tectonics of the Southern Central Andes. Springer, Berlin, pp. 121- 139. Shimamoto, T., 1989. The origin of S-C mylonites and a new fault-zone model. J. Struct. Geol., 11: 51-64. Steiger, R.H. and Jlger, E., 1977. Subcommission on geochronology: Convention on the use of decay constants in geo- and cosmochronology. Earth Planet. Sci. Lett., 36: 359-362. Stormer, J.C., 1975. A practical two feldspar thermometer. Am. Mineral., 60: 667-674. Treloar, P.J., Rex, D.C., Guise, P.G., Coward, M.P., Searle, M.P., Windley, B.F., Petterson, M.G., Jan, M.Q. and Luff, I.W., 1989. K-Ar and Ar-Ar geochronolopgy of the Himalayan collision in NW Pakistan: constraints on the timing of suturing, deformation, metamorphism and uplift. Tectonics, 8: 881-909. Tsemr, MC. and Carter, N.L., 1987. Upper limits of power law creep of rocks. Tectonophysics, 136: l-26. Tullis, J., 1990. Experimental studies of deformation mechanisms and microstructures in quartzo-feldspathic rocks. In: D.J. Bar ber and P.G. Meredith (Editors), Deformation Processes in
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Minerals, Ceramics and Rocks. (Mineralogical Society Series.) Unwin-Hyman, London, pp. 190-227. Twiss, R.J., 1977. Theory and applicability of a recrystallized grain size paleopiezometer. Pure Appl. Geophys., 115: 227244. Van Blanckenburg, F., Villa, I.M., Baur, H., Morteani, G. and Steiger, R.H., 1989. Time calibration of a PT-path from the Western Tauem Window, Eastern Alps: the problem of closure temperatures. Cotmib. Mineral. Petrol., 10 1: 1- 11. Wheeler, J., 1987. The significance of grain-scale stresses in the kinetics of metamorphism. Co&b. Mineral. Petrol., 97: 397404. Wintsch, R.P. and Andrews, M.S., 1988. Deformation-induced growth of sillimanite: “stress” minerals revisited. J. Geol., 96: 143-161. Wintsch, R.P. and Dunning, J., 1985. The effect of dislocation density on the aqueous solubility of quartz and some geologic implications: a theoretical approach. J. Geophys. Res., 9O(B5): 3649-3657. Woodcock, N.H., 1986. The role of strike-slip fault systems at plate boundaries. Philos. Trans. R. Sot. London, Ser. A, 317: 13-29. Zonenshayn, L.P., Savostin, L.A. and Sedov, A.P., 1984. Global paleogecdynamic reconstructions for the last 160 million years. Geotectonics, 18: 181-195.
Tectonophysics 250 (1995) 61-87
JOAr/ “Ar and Rb-Sr analyses from ductile shear zones from the Atacama Fault Zone, northern Chile: the age of deformation Ekkehard Scheuber a,*, Konrad Hammerschmidt
b, Hans Friedrichsen
b
’ lnstitutfir
Geologic. Geophysik und Geoinformatik, Freie Uniuersitiit Berlin, Malteserstrasse 74.100, D-12249 Berlin, Germany h Institutfir Mineralogie, FR Geochemie, Freie Universittit Berlin, Boltzmannstrasse 18-20, D-14195 Berlin, Germany
Received 19 July 1994; accepted 6 April 1995
Abstract The influence of deformation on the K-Ar and the Rb-Sr isotope system is investigated. It is assumed that, due to the diffusion processes involved, deformation has a similar effect on isotopic equilibrium as has temperature. In order to examine the influence of deformation on the K-Ar and the Rb-Sr isotope systems two shear zones from the Atacama Fault Zone (AFZ), situated in the north Chilean Coastal Cordillera, have been investigated. The AFZ, which was active as a sinistral strike-slip fault during the Mesozoic, has two sets of shear zones, one formed under amphibolite (SZl), one under greenschist facies conditions (SZ2). Rb-Sr and 40Ar/39Ar age determinations were conducted on samples from cross sections of each set. In SZl the homblendes and biotites from a weakly deformed sample reveal cooling ages of 153-152 and 150 f 1 Ma, respectively. Biotite from the center of the shear zone of SZl gave an isochron of 143.9 + 0.3 Ma (MSWD = 0.04) which is interpreted as the age of deformation which produced resetting of the mineral system. In SZ2 homblendes yielded “OAr/ 39Ar plateau (cooling) ages of - 138 Ma. Biotites from undeformed samples gave Rb-Sr and “OAr/ 39Ar total degassing ages of 130 f 1 Ma, whereas biotite from the mylonitic rocks yielded 126- 125 Ma which dates the time of deformation. Sr isotope homogenization occurred in the mylonitic rocks, and is most likely a result of deformation. The formation of SZl can be correlated to the Araucanian (= Nevadan) phase. The deformation in SZ2 is related to the onset of uplift and cooling of the Coastal Cordilleran magmatic arc.
1. Introduction Geochronological methods are a powerful tool for the dating of semiductile and ductile deformational structures (e.g., Kligfield et al., 1986; Hubbard and Harrison, 1989; Treloar et al., 1989; Getty and Gromet, 1992; Leloup et al., 1993). However, dating deformational events often meets the problem of how to distinguish ages due to regional cooling from
* Corresponding author. 004O- 195 l/95/$09.50 0 1995 Elsevier Science B.V. All rights reserved LSD1 0040- I95 1(95)00044-5
ages due to deformation because ductile deformation of quartz-feldspar rocks is only possible in nature under elevated temperatures (greenschist facies or higher). Cooling ages are associated with the closing of isotopic systems of a mineral by passing through its closure temperature CT,>. In the relation between the age of a mineral, recrystallized during deformation, and passing the T, two cases can be distinguished: (1) When deformation takes place before the mineral falls short of T,, the age of this mineral is a pure cooling age which only gives a minimum age of deformation. (2) When deformation takes place after passing T,, it is possible to determine the
62
E. Scheuber et al./ Tectonophysics 250 (1995161-87
age of deformational events if the mineral under question recrystallizes and attains complete isotopic equilibrium. In literature it is assumed that the critical role in age resetting is played by mass transfer due to diffusion (Dodson, 1979). Diffusion is facilitated not only by elevated temperatures, but also by differential stresses and strain energy (e.g., Wintsch and Dunning, 1985; Wheeler, 1987) which have a great influence on chemical potentials and, for example, are driving forces for metamorphic reactions (Wintsch and Andrews, 1988; Klaper, 19901. In this paper we discuss the influence of deformation on the Rb-Sr and K-Ar systems of mylonites from the Atacama Fault Zone (AFZ), northern Chile; here, two sets of ductile shear zones occur which originated under amphibolite facies and greenschist facies conditions, respectively (Fig. 1). Age determinations of biotite and hornblende have been carried out in cross sections from the undeformed protolith into the center of shear zones crosscutting plutonic rocks. It will be shown that deformation produced a resetting of Rb-Sr and K-Ar mineral ages, and that this resetting was accompanied by homogenization of the Rb-Sr whole-rock system at least on the meter scale. However, resetting and homogenization was only detectable in minerals of sheared rocks where deformation started after cooling of the undeformed protolith below T,; the age of high-grade deformations in an amphibolite facies shear zone gave only the age of cooling after deformation. I .I. Plastic deformation
and closure temperatures
The theory of closure temperatures (Dodson, 1973, 1976, 1979) is based on the assumption of a firstorder loss or of volume diffusion of the radiogenic nuclides (e.g., Chermiak and Watson, 1992). T,, which is defined as the temperature (T,) of the system at its calculated age, is given by: E DO = In (~72 i i RT,
Fig. 1. Map of the north Chilean Coastal Cordillera showing the location of shear zones along the Atacama Fault Zone, and the investieated cross sections SZl and SZ2.
E. Scheuber et al./
Tectonophysics 250 (1995) 61-87
where E is the activation energy, R the gas constant (R = L X k, L is the Avogadro constant, k is the Boltzmann constant), D, the frequency factor (diffusion coefficient at infinitesimally high temperature), a the radius of the mineral grain, (Y a geometric factor and r a cooling-time constant (r = - RTcz/[ E(dT/dt)]} which considers the cooling rate (dT/dt) of a particular area (T is temperature and r is time). In plastic deformation of silicate rocks under natural conditions the most common deformation processes are diffusional flow and dislocation creep. In both processes the strain rate is controlled by the applied stress and by the diffusion of vacancies, ions or atoms (e.g., Poirier, 1985). Stress (a) and strain rate (+) are related by the linear form (cf. Frost and Ashby, 1982; Poirier, 1985): D eff in the case of diffusion form:
creep, and by the power-law
&=A( gneXP( - $,) in the case of dislocation creep (power-law climbplus-glide sensu Frost and Ashby, 19821, where (Y is a numerical factor depending on the shape of the grain, R is the atomic or molecular volume, T is the temperature, a is the grain size, De,, is the effective diffusion coefficient, A is the power-law creep constant, p is the shear modulus and n is the stress exponent (stress sensitivity). Dislocation creep causes lattice defects to migrate through crystals under an applied stress to form suband new-grain boundaries by either migration recrystallization or rotation recrystallization (cf. Drury and Urai, 1990). Migration recrystallization is in particular expected to support chemical and/or isotopic exchange (Getty and Gromet, 1992), as in this process regions of high dislocation density are repeatedly replaced by dislocation-free material as mobile grain boundaries migrate through crystals. Diffusion creep is the predominant deformation process at very small grain sizes ( < 10 pm for quartz, Handy, 1989); it is also expected to support chemical and isotopic exchange as it results from stress-directed
63
bulk vacancy diffusion (Tsenn and Carter, 1987) along grain boundaries or through the lattice. Diffusion creep includes pressure solution which is strain and stress controlled. For example, Wintsch and Dunning (1985) have shown that the solubility of quartz increases if tangled dislocations are present (‘‘strain solution’ ‘). Grain-scale differential stresses induce pressure solution and are a driving force for chemical exchange (Wheeler, 1987; Wintsch and Andrews, 1988). Although at present no theory is available which links the theories of closure temperature and plastic deformation, a comparison of both suggests that, due to the diffusion processes involved, both strain rate and temperature have a similar effect on closure or opening of isotope systems. This means that an increasing strain rate should lead to age resetting in an environment of uniform temperature, even below T, and although resetting may be enhanced by shear heating. Therefore, if a gradient in strain rate is observed, age resetting should preferably affect those rocks which accommodate higher strain rates. This means that, in inhomogeneous materials where the strain rate concentrates on the weaker parts, age resetting should preferably affect the weaker material. Furthermore, if the strain rate drop after deformation is rapid, the relevant isotope systems of all minerals that recrystallized dynamically will be “frozen in”. This is similar to the preservation of microstructures which also depends on the rate of strain rate drop (Knipe, 1989). On the other hand, a slow decrease in stress/strain rate should have the same effect as if the temperature remains high: the deformational microstructures should be obliterated by annealing; in this case the minerals are expected to show a normal cooling-like pattern, and the minerals show different ages. The finite strain is also expected to support age resetting as, with increasing finite strain, the number of grains undergoing dynamic recrystallization increases. This is similar to the time dependence of heating on isotopic exchange: the duration of heating controls the degree of isotopic exchange. A very important effect of dynamic recrystallization which is expected to support isotopic exchange is the deformation-induced grain size reduction. The grain size-stress relation (e.g., Twiss, 19771, which has been established theoretically for migration re-
64
E. Scheuber et al./ Tectonophysics
crystallization by Derby (1990), says that the smaller the recrystallized grain size the higher the applied stress. Smaller grain sizes, on the other hand, reduce the closure temperature (Eq. 1). Therefore, it is possible to observe large but older grains whose age values survive in a matrix of younger but smaller grains. From these considerations it is concluded that the age of a deformational event can be dated if strain rate, strain and/or deformation temperature are sufficient to produce a resetting of the investigated isotope systems. To distinguish an age due to deformation from an age due to regional cooling the following two criteria should be fulfilled: (1) deformation should have started below the closure temperature of the host rock; and (2) deformation-induced resetting should be complete. Criterion (1) is fulfilled if the same mineral species is younger in the tectonite than in the host rock. Criterion (2) is fulfilled if deformed or recrystallized minerals show isotopic equilibrium with other minerals.
2. Geological
setting
During the Jurassic and Early Cretaceous a continental magmatic arc developed in the north Chilean Coastal Cordillera. It consists of large quantities of basaltic to andesitic lavas (La Negra Formation, Garcia, 1967) and extended batholiths of gabbroic to granodioritic composition. Magmatic activity lasted from the Early Jurassic to the Early Cretaceous, as shown by isotope age data (200-I 10 Ma; e.g., Halpem, 1978; Diaz et al., 1985; Maksaev et al., 1988; HervC and Marinovic, 1989). The tectonic setting of the Jurassic-Early Cretaceous arc can be described as extensional to transtensional (Scheuber and Reutter, 1992; Grocott et al., 1994). Its major structure is the orogen-parallel Atacama Fault Zone (AFZ) which can be traced for more than a thousand kilometers from _ 20” to N 30%. Scheuber and Andriessen (1990) gave a detailed description of AFZ structures south of Antofagasta (23”45’-25%) where two sets of ductile shear zones developed (Fig. l), one formed under amphibolite facies conditions and one under greenschist facies conditions (called Jurassic and Early
250 (1995) 61-87
Cretaceous shear zones, respectively, by Scheuber and Andriessen, 1990). In the Early Cretaceous shear zones the vertical foliation planes generally trend N20”E and contain subhorizontal stretching lineations. The sense of shear is uniformly sinistral, which is consistent with highly oblique ( > 45”) subduction towards te southeast during the Late Jurassic-Early Cretaceous (Zonenshayn et al., 1984, Jaillard et al., 1990); the AFZ of that time was thus interpreted as a trench-linked strike-slip fault sensu Woodcock (1986). In the Jurassic shear zones the general trend of the foliation planes is also N20”E. However, kinematic indicators are less well developed and rarer than in the Early Cretaceous shear zones probably due to thermal or high-temperature tectonic overprinting. The kinematic pattern is complicated and not yet fully understood. Structures indicative of sinistral strike slip were overprinted and partly obscured by vertical displacements in some places (Gonzalez, 1993). Scheuber and Andriessen (1990) have published evidence for sinistral strike-slip movements. Structures due to vertical displacements have been described by Brown et al. (1993) from the Coastal Cordillera between 25 and 27%. These authors attribute such movements to the subsidence of a marine backarc basin to the east of the magmatic arc. In some places an overprinting of the structures due to strike-slip movements by dip-slip normal movements can be observed. One example is illustrated in Fig. 2. At the southern contact of the Pluton Cristales (Rb-Sr whole-rock age 145 + 10 Ma, Sr, 0.70326, Her& and Marinovic, 1989) the overall N-S trend of the foliation planes is overprinted by a second generation of planes running parallel to the contact of the Pluton Cristales. Near the contact the older planes are completely obliterated, whereas at a greater distance from the contact both planes can be distinguished. The older foliations contain subhorizontal stretching lineations, and are thus due to strike-slip movements; on the other hand, the younger foliations contain down-dip lineations, and kinematic indicators reveal subsidence of the Pluton Cristales with respect to its host rock. The older (strike-slip) deformation led to the formation of a recrystallized hornblende-plagioclase fabric, whereas the younger foliation planes are defined by the preferred orientation of biotite.
65
E. Scheuber et al. / Tectonophysics 250 (I 995) 61-87 70°25’
70030’ w
24’=05’
v;,x\,x x v \\ x--x-
x /_
,”
x x
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0
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5 km 1
2
3
4
2. Geological map of the Coastal Cordillera near Cerro Cristales showing the location of the investigated shear zone SZI. I = Jurassic intrusive rocks (Paranal Unit, - 170 Ma); 2 = Jurassic volcanic rocks; 3 = Pluton Cristales (- 143 Ma); 4 = trend of foliation planes; sources: own mapping and Gonzalez (1990).
Fig.
3. Lithology and structures 3.1. Amphibolite
facies shear zone (SZI)
SZl, approximately 15 m wide, is located N 50 m south of the contact to the Pluton Cristales (Figs. 1 and 2). Table 1 shows the locations, modal compositions, and structural features of the investigated samples. The vertical foliation planes of SZl strike N60”E, the older planes cannot be observed macroscopically. Due to increasing strain, the spacing of
foliation planes becomes narrower approaching the shear zone center. The shear zone is developed in a hornblende gabbro of the Paranal Unit (Rb-Sr whole-rock age 170 f 28 Ma, Sr, 0.70315, HervC and Marinovic, 1989) containing some biotite, orthoand clinopyroxene (sample 87-150). In the shear zone the gabbro is altered to an amphibolite gneiss (samples 87-153 10 m north of the shear zone center, 87-151 shear zone center, Fig. 3) consisting mainly of plagioclase and magnesio-hornblende. The Ancontent of plagioclase decreases with increasing strain
66
E. Scheuber et al. / Tectonophysics
from N 45 mole% (87-153) to 32 mole% (87-151) (microprobe analyses). In the shear zone center some newly crystallized orthopyroxene is present (Fig. 3a) indicating upper amphibolite facies metamorphic conditions (Bucher and Frey, 1994). The biotite content increases from N 1% in the undeformed gabbro to 7-9% in the shear zone probably as a result of a better permeability for fluids due to the smaller grain size in the shear zone. The growth relationships of biotite reflect both deformations which affected the SZl rocks; the following types of biotite grains can be distinguished: (1) old grains of magmatic origin; (2) flakes which have grown across the recrystallized plagioclase-hornblende fabric without preferred orientation, and which are frequently bent (Fig. 3b);
250 (I 995) 61-87
and (3) partly recrystallized grains with straight grain boundaries oriented parallel to the foliation (Fig. 3~). Old grains (1) form only a minor fraction of biotite, their occurrence being restricted to the undeformed (87-150) and the moderately deformed samples (87153); grains of type (2) also occur mainly in the moderately deformed sample, to a lesser extent also in the shear zone center; biotite grains of type (3) dominate in the shear zone center but they are also present in the shear zone margin. From the growth relationship between biotite, plagioclase and homblende the following run of events can be inferred: biotite of type (2) grew statically after the first deformation which produced the recrystallized plagioclase-hornblende fabric. During the second defor-
Table 1 Location, modal composition, and structural features of the analyzed samples: Modal composition (vol.%) of rock-forming minerals, the amount of recrystallized (recr.) and magmatic (magm.) grains; average grain size of matrix grains; and ellipticity of the finite strain ellipse in X2 sections [ R(XZ)]. In SZl amd SZ2, the samples are grouped from left to rigth, respectively, with decreasing distance from the shear zone center
SZl 87-150 70025.5’W 2.4011.5'S
SZ2
87-153
87.151
70'25.5'W 24011.5'S
70025.5'W 24or1.5s
85-58 70*15.5'w !4040.
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88-26
88-1 a
70~17.3'W 24o38.1'S
70'19.3'W 24O37.I'S
88.lc 70~1.9.5'W 24O37.6'S
88-10
88-16
85.546
70~1.9.8'W 24O37.3'S
70"20.6'W 24038.3'S
70°20.6% 24O38.3'S
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7.5 0.5 8.0
0.0 0.0 0.0
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21.6 0.0 21.6
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17.5 6.2 23.7
4.3 20.5 24.8
18.6 19.6
0.1 30.3 30.4
0.0 13.8 13.8
56.9 6.5 63.4
35.6 27.9 63.7
3.5 61.7 65.2
53.3 0.0 53.3
58.4 0.0 58.4
51.4 0.0 51.4
46.4 1.7 48.0
47.9 3.3 51.3
41.4 3.3 44.4
0.9 73.2 74.1
0.8 0.0 6.6
0.9 0.0 0.9
0.0 0.0 0.0
10.7 0.0 10.7
6.6 0.0 6.6
12.6 0.0 12.6
8.4 3.0 11.4
5.2 1.1 6.4
9.7 6.1 15.7
0.0 0.0 0.0
23.7 0.0 23.7
9.2 14.0 22.2
0.2 22.9 23.1
5.9 0.0 5.9
6.8 0.0 6.8
5.5 0.0 5.5
5.9 0.5 6.4
5.2 1.1 6.3
1.6 0.1 1.7
0.0 0.0 0.0
1.2 0.0 1.2
3.3 5.6 8.9
0.2 6.6 6.8
7.1 0.0 7.1
5.6 0.0 5.8
6.0 0.0 6.0
2.5 6.7 9.2
2.7 11.6 14.3
1.6 4.7 6.3
0.0 0.0 0.0
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0.105
1.4 _.. -1
a Igneous pyroxenes. b Metamorphic pyroxenes. ’ Porphyroclasts.
._. -1
-1
>>50
61
E. Scheuber et al. / Tectonophysics 250 (1995) 61-87
The rocks of SZl were deformed within a fully ductile rheology without any ductility contrast between each mineral species. This is revealed by a steady increase in the amount of recrystallized grains of all mineral species from the margin into the center of the shear zone (Table 1; Fig. 41, and by the observation that there is no grain size reduction around old grains which thus did not operate as rigid particles but rather are remainders due to incomplete recrystallization. In the shear zone center recrystallization is nearly complete, whereas in the sample from the margin 40% of old grains are present. Postdeformational annealing in the rocks of SZl is revealed by a granoblastic fabric of the matrix where many grains have straight grain boundaries and 120” triple points are frequently found in plagioclase domains. These observations indicate that the peakstress event is not preserved in the microstructures of the rocks (Knipe, 1989; Prior et al., 1990) due to a relatively slow regional cooling and/or a slow decrease of stress after deformation. In the annealed fabric of the rocks from SZl it is not possible to determine the finite strain. 3.2. Greenschist facies shear zone (SZ2) SZ2 is located in a shear zone which runs through a medium-grained granodiorite of the Remiendos Unit (Her& and Marinovic, 1989: Rb-Sr biotite age
l-
Fig. 3. Microphotographs of sheared rocks of SZI. (a) Newly (Opx) in sample (87-1.51) from the crystallized orthopyroxene shear zone center. (b) Biotite flakes (B) which grew across the recrystallized, pophyroclastic fabric of plagioclase and hornblende (sample 87-153, 10 m north of shear zone center). (c) in the shear zone center (sample 87-151) biotite is recrystallized dynamically.
g
0.8 . .
f 8 2
0.6 .a
g r h g
?? 0.4 . . 0.2 a* 0*.. 0
A :
0.2
.
: 0.4
.
: 0.6
AA : 0.8
1
(qtz+bi)recr/qtz+bi)tot mation these grains were deformed and recrystallized to form type (3). Thus, recrystallized biotite (type 3) should be younger than grains of type (21, and of course also younger than igneous biotite of type (1).
Fig. 4. Diagram showing the degree of recrystallization (recr = recrystallized: tot = whole rock) of incompetent minerals (qtz = quartz; bi = biotite) and competent minerals (fsp = feldspar; hb = hornblende); squares: SZl, triangles: SZ2.
68
E. Scheuber et al./ Tecronophysics 2.50 (I9951 61-87
13 1 i- 1 Ma). The exposed width of the mylonitic zone is about 2 km. Samples have been collected from the undeformed granodiorite (samples 85-58 and 88-26), and across a section through the shear zone (samples 88-la, c, e, g and 8554b) (Table 1; Fig. 2). A NNE-striking subvertical foliation develops N 1.9 km east of the shear zone center. Due to increasing finite strain the spacing of the foliation planes becomes narrower and the subhorizontal stretching lineations become more distinct to the shear zone center. Except for ultramylonites from the center of SZ2 the sheared rocks are medium- to fine-grained S-C mylonites (Lister and Snoke, 1984) showing a sinistral sense of shear. The granodiorites and the mylonites are composed of plagioclase (An _ so), K-feldspar and quartz. From the composition of recrystallized feldspar in the mylonites (albite in plagioclase vs. albite in alkalifeldspar geothermometer of Stormer, 1975) metamorphic temperatures can be roughly estimated as 400-500°C. Mafic minerals in the granodiorite and the mylonites are mainly magnesio-hornblende and biotite. In the granodiorites hornblende occurs as euhedral crystals, sometimes with cores of clinopyroxene. In the mylonites hornblende forms patchy porphyroclasts surrounded by rims of very finegrained biotite and actinolite often developed as trails extending into C-planes. Very small biotite flakes also grew along cracks in the interior of hornblende porphyroclasts (Fig. 5a). Most biotite occurs as rounded, partly recrystallized aggregates (former igneous crystals), a minor amount grew along
Fig. 5. Microphotographs of mylonites of SZ2. (a) Hornblende porphyrcclast (H) in center of SZ2 (sample 88-lg). Biotite (b) growth occurred on the rim, on cleavage planes and in a shear plane; on the rim of the hornblende porphyroclast also some recrystallization of hornblende (h) took place. The shear plane on the upper left side is rather sharp, according to Shimamoto ( 1989) indicating semiductile deformation related to some extent to seismic slip events. (b) Two contrasting sites of biotite recrystallization in a granodioritic mylonite &rnple 88-le). B = large, former igneous grains; (I = small flakes in shear plane. The plagioclase porphyroclasts (at the upper margin of the photograph) shows dynamic recrystallization on the rim along the shear plane. (c) Diopside porhyroclasts (D) in a very fine grained dynamically recrystallized matrix of mainly plagioclase (P), some coarser grained quartz is also present in the matrix, e.g., on the right side above the scale bar.
shear planes (Fig. 5b). The major part of the matrix of the mylonites consists of quartz. The ultramylonites (sample 85-54b) form bands, up to 50 m wide, which run through the “normal” mylonites parallel to the foliation. In these rocks no S-C fabric but extensional crenulation cleavages are developed. The ultramylonites consist mainly of intermediate plagioclase, quartz, diopside, epidote and sphene (Fig. 5~). Plagioclase and, to a minor extent,
E. Scheuber et al. / Tectonophysics 250 11995) 61-87
quartz form the ductile matrix; diopside, epidote and sphene are porphyroclasts. Compared with the “normal” mylonites in the ultramylonites a complete loss of K-feldspar, biotite, hornblende and Fe-ore minerals has taken place which, together with the very fine-grained matrix of the ultramylonites with an average matrix-grain size of N 4 pm, indicates that diffusional mass transfer played an important role in the formation of this rock. In contrast to SZl the incompetent minerals of SZ2 mylonites (quartz and biotite) were much more strongly affected by dynamic recrystallization than the competent minerals (feldspar and hornblende; Table 1; Fig. 4). This ductility contrast indicates that these rocks deformed in a semiductile manner as suggested by Shimamoto (1989) for S-C mylonites. However, the competent minerals are not completely free of dynamic recrystallization, as the amount of dynamic recrystallization of these minerals increases from 5% (88-1~) to 9% (SS-lg) from the margin into the shear zone center. In the ultramylonitic sample (8554b) all plagioclase has recrystallized dynamically. In the S-C-mylonites quartz is the weakest phase, deformed mainly by dislocation creep (indicated by lattice preferred orientations, Scheuber and Andriessen, 1990); quartz is weaker than biotite which forms elongate aggregates (fishes). Again the ultramylonites are an exception in that quartz forms elongate lenses in a plagioclase matrix indicating that plagioclase is weaker than quartz; this is consistent with the findings of experiments reported by Dell’Angelo and Tullis (1982) and Tullis (1990). They described a very distinct strain weakening of feldspar due to dynamic recrystallization by straininduced grain boundary migration. The SZl mylonites contain 19-31% of quartz (exception: ultramylonites 13%). The size of recrystallized new quartz grains is N 45 pm; from this size a flow stress in the order of 40-60 MPa can be calculated (Scheuber, 1994) using the paleopiezometers of Twiss (1977). The quartz content of the mylonites is in the range where a transition takes place from a framework-supported rheology to a matrix-supported rheology (e.g., Jordan, 1988; Handy, 1990). This is consistent with the microstructures of the mylonites: In those samples with a high quartz content (e.g., sample SS-lg: 3 1% qtz) the quartz lenses are more interconnected in matrix form
69
than in the quartz-poor rocks. For example, in sample 88-le (19% quartz) quartz forms elongated lenses oriented parallel to the foliation; however, these lenses are often isolated and do not form an interconnected matrix. Thus, the quartz-poorer samples were stronger during deformation and, consequently, accommodated a smaller fraction of the bulk strain rate of the shear zone than those samples richer in quartz. The finite strain (X/Z ellipticity R) increases from R - 2 (sample 88-1~) to R - 4 (SS-lg) measured using the normalized Fry method (Table 1; Erslev, 1988). The highest strain was accommodated by the ultramylonites with R B- 50 (R,/cp’ method in quartz lenses). However, as most measured particles are disrupted it is not possible to determine which fragments are part of one pre-deformational clast and thus the measured R, value is generally too low. The high strain in the ultramylonites is shown by apatite grains with R, > 1000, in X§ions the length of these grains exceeds 1 cm, whereas their width is only N 10 pm. There are no indications of annealing of the mylonitic fabrics in SZ2. This is inferred from the observation that the grain size of dynamically recrystallized new grains is smallest at sites of stress concentration, e.g., around porphyroclasts or in Cplanes. The varying sizes of new grains in the mylonites of SZ2 indicates that due to a rapid stress drop the peak stress event has “frozen in” the size of the new grains (Knipe, 1989).
4. Results of age determinations Rb-Sr and “OAr/ 39Ar data are listed in Tables 2-4. In Table 5 we summarize the Rb-Sr ages and the “OAr/ 39Ar total degassing ages as well as the plateau ages in the order of decreasing distances from the shear zone centers. In the column “Geological mean age” the mean values of geologically significant @Ar/ 39Ar and Rb-Sr ages are shown. Rb-Sr isochrons (Figs. 6 and 7) were calculated only if the data of at least three mineral components of one hand specimen are available (Table 4). @Ar/ 39Ar total degassing ages were deduced from summing up the data of the incremental heating steps (Table 4). Plateau ages (McDougall and Harrison, 1988) are considered when three or more successive incremen-
70
E. Scheuber et al./ Tectonophysics
tal degassing steps give, within the 2u error limits, the same age value and when the trapped “OAr/ 36Ar ratios are identical to the atmospheric value (295.5). Table Results
250 (1995) 61-87
The trapped 4oAr/ 36Ar ratio is given as the intercept of a line defined by these plateau age steps in a discrimination diagram of 40Ar/ ‘(jAr vs. 39Ar/ 36Ar.
2 of the Rb-Sr
analyses
Sample SZl 87-150
87-153
for whole rock and minerals
material
88-26
88-1 a
88-lc
88-le
88-19
85.54b
Sr [nom1
87Rb/86Sr
20
87Sr/ssSr
16.4 14.1 6.57
471.4 705.3 88.4
0.1008 0.0580 0.2771
0.703496 0.703402 0.704172
* * *
120 45 30
0.25-0.5 0.25-0.5 0.25-0.5
15.3 4.35 5.48 349.8
461.6 786.6 65.5 37.3
0.0961 0.0160 0.2423 27.32
0.703354 0.703258 0.704019 0.761504
f * f +
30 30 180 35
Pig hb bi
0.25-0.4 0.12-0.25 0.25-0.4
20.6 2.51 6.43 307.4
396.7 569.3 31.6 11.6
0.1503 0.0127 0.5902 78.01
0.703795 0.703519 0.704724 0.863298
+ f + r
35 40 35 50
WR LM bi
0.13-0.4 0.13-0.4
55.5 24.6 303.8
360.5 341.7 12.1
0.4455 0.2083 73.33
0.704265 0.703729 0.837958
+ * r
50 50 120
WR LM hb bl
0.25-0.7 0.25-0.7 0.5-0.7
48.4 27.1 3.18 327.2
407.4 416.1 25.9 7.6
0.3424 0.1884 0.3571 126.8
0.703934 0.703681 0.703932 0.936388
+ + * *
80 50 50 120
WR LM bi
0.25-0.5 0.25-0.5
86.6 53.0 556.8
389.3 381.8 20.8
0.6420 0.4004 78.55
0.704652 0.704176 0.847446
+ f -f
70 35 85
WR LM hb bl
0.12-0.25 0.181-0.25 0.12-0.25
78.8 45.6 31.6 552.2
405.8 437.8 27.4 20.9
0.5615 0.3018 3.338 77.50
0.704484 0.703971 0.709735 0.841503
k * * k
150 30 45 50
WR LM bl
0.12-0.25 0.12-0.25
59.7 20.6 439.2
420.3 437 2 24.6
0.4104 0.1346 52.06
0.704236 0.703692 0.799226
+ * +
140 45 70
73.8 61.5 148.0 2.85
403.7 429.2 68.3 391.4
0.5295 0.4137 6.267 0.0210
0.704477 0.704234 0.714746 0.703551
* + + t
50 60 25 70
WR hb bi
sz2 85-58
Rb lppml
0.25-0.5 0.25-0.5
WR LM hb
PM
87-151
grain size (mm)
WR
WR LM hb( + bil WR
0.12-0.25 0.12-0.25
WR = whole rock; plg = plagioclase; LM = light minerals (plagioclase, quartz, K-feldspar); hb = hornblende; bi = biotitc; the grain size (mm) of the analyzed minerals is also given; errors refer to the last digits. Analytical methods: Aliquots (0.1 g) of whole rocks (3-5 kg), plagioclase or light minerals, hornblende and biotite were dissolved in a mixture of hydrofluoric and nitric acid under a pure nitrogen atmosphere. Sr and Rb concentrations were determined using a highly enriched 84Q-spike,and a *‘Rb-spike solution adding to aliquots of the sample solution. Rb and Sr were separated using ion exchange in resin columns. Approximately 0.1 g samples were processed, and the blanks never exceeded 0.50 ng Sr and 0.15 ng Rb. The precision on Sr and Rb concentrations is estimated as 0.1 and OS%, respectively (lo error). The isotope compositions of the samples were measured on a MAT 261 (Finnigan) thermionic mass spectrometer operating in the static mode. The relative transmission of the ion beam between the different faraday cup positions was in the order of 3 X IO-‘. Routinely, the measured NBS 987 standard gave s7Sr/ s6Sr = 0.71022 + 4 (2 u~,,,~) normalized to a value of s6Sr/ ‘*Sr = 0.1194.
E. Scheuber et al./ Tectonophysics
71
250 (1995) 61-87
Sr (e.g., van Blanckenburg et al., 1989). Any deviation from this behaviour is the result of a geological or an artificial disturbance. Such disturbances can be recognized in a staircase age spectrum with the “OAr/ 3gAr method. In the case of biotite there are at least four possibilities of explaining such staircase age spectra: (1) Thermal overprinting which leads to a partial Ar loss (Harrison, 1983); when the thermal pulse is moderate the Rb-Sr biotite system is not affected and as a consequence the 4oAr/ 3gAr total degassing age is lower than the Rb-Sr age. (2) Radiogenic Ar loss during very slow cooling (< l”C/Ma) which also generates staircase age patterns (McDougall and Harrison, 1988). In this case the closure of a system is a temperature range where at high temperatures only a small amount of radiogenie Ar is stored within the mineral and by far the largest amount of Ar is lost by diffusion, at the lower end of this range nearly all radiogenic Ar is kept in the mineral. Chemical bonding of Sr is stronger and therefore Rb-Sr ages are higher than the K-Ar c40Ar/ 39Ar total degassing) ages. (3) Vacuum degassing effects are thought to change the “OAr/ 39Ar distribution. Recent results
Two types of age spectra can be distinguished (Table 5; Figs. 8 and 91: flat plateau ages across a wide range (> 70%) of 3gAr released (e.g., sample 87-150 hornblende, Fig. 8), and staircase patterns sometimes with a plateau-like part of 20-55% of 3gAr released. In relation to the Rb-Sr ages two kinds of plateaus can be distinguished: in SZl samples the plateau age values agree with the Rb-Sr ages and the “OAr/ 3gAr total degassing ages are younger than the Rb-Sr ages (e.g., sample 87-153 biotite, Fig. 8); by contrast, the total degassing ages of SZ2 samples are equal to the Rb-Sr ages whereas the plateau-like parts reveal higher age values (e.g., sample 88-la biotite, Fig. 9). Thus, we have to interpret these staircase age patterns. In thermally undisturbed samples K-Ar biotite ages which are identical to 40Ar/ 3gAr total degassing ages agree with the Rb-Sr ages because both systems shared the same closure temperature of about 300°C (Purdy and Jager, 1976; van Blanckenburg et al., 1989; Hurford et al., 1989). Recently the closure temperature of hornblende K-Ar was revised by Dallmeyer (1978) as 480°C and by Harrison (1981) as 53Or$,“C. Therefore, in a steadily cooling environment we can observe the following age sequence: hornblende K-Ar > biotite K-Ar, Rb-
Table 3 Rb-Sr results
Sample
lsochron I Reference line
SZI 87-153
MSWD
Age
_
e’SrleeSri
20
0.703149
f
25
0.22 0.04
0.703486 0.703490
f +
32 30
i 0.4
2.03
0.703397
f
70
129.3
i
0.3
0.38
0.703326
f
45
WA-LM-bi
129.0
i
0.3
0.65
0.703448
+
60
88-1~
WR-LM-hb WR-LM-bi
133.6 125.3
f 1.2 f 0.3
0.07 0.21
0.703390 0.703434
f f
20 20
88-le
WR-LM-bi
129.4
f 0.3
0.24
0.703448
f
45
88-lg 88-l g + 85-54b
WR-LM-(hb + bi) WR-LM-(hb + bi)WR
126.1 126.1
+ 0.6 f 0.6
0.79 0.39
0.703515 0.703515
k *
40 36
WR-bi
150.3
f 0.3
WR-pig-hb WR-pig-bi
140.5 143.9
f 7.9 f 0.3
85-58
WR-LM-bi
129.1
88-26
WR-LM- bi
B&la
87-151
522
WR = whole rock; pig = plagioclase;
LM = light minerals;
hb = hornblende;
bi = biotite. 2~ error values refer to the last digits.
f f zt
4869 8413 2375
f
4691
i
f
4000
2310 4516
1200 4000
*
*
*
+
f
*
f
4000
1765
1161
750 800
7194
2110
4800 3000
1785
2100
434
*
f f f +
3697
2590
4170
12270
3145
5675
5745
10050
3960
3365
8640
1905
2891
975
990
1005
1020
1040
1070
1120
1130
1180
1240
1320
Total
+
r
f
f
*
*
*
*
+
*
1283 2059
6561
11000 1500
2388
1500 4000
2922
7426
4284
4237
2500
10000
5000
4000
2316
3180
3200 9287
1917
1100
1400
2576
3500
10000
3987
5000
29.7 305
15 100
+
i
+
f
f
f
f
f
f
f
f
*
*
f
*
*
.
0.1540
1100
2700
8600
1000
0.1192
0.1162
0.1140
0.1159
0.1131
1800
0.1285 0.1114
0.1175
3100 3600 7000
0.1115
0.1050
0.1003
0.1118
1100
9000
2400
810
0.3752
1.165
2300
1.323
0.4655
45 3000
1.1
Le 0.26-0.5
950 Sli
0.1248
130
grain
0.1721
0.1504
0.1353
0.1319
0.1292
0.1421
2800
3000
3000
850
550
500
0.2650 0.1612
0.2802
1400 5600
0.6460
3000 4000
1.082
35
*
f
It
*
f
k
2
*
*
*
i
+
k
*
at
- 00002 0.0002
313.2
2.0
24.0
0.0003
15.66
9.17
14.67
145.9
20.2
14.19
12.13
12.78
5.12
3.40
3.43
0.81
88 * 10
415.8
56.2
29.8
I [
174.3
58.3
62.4
24.0
15.93
4.12
3.91
1.31
3.36
4.63
7.28
I.5 mm,
Chile)
zt
i
*
*
f
*
+
f
f
*
*
*
*
*
*
f
6.9
1.1
1.2
1.3
0.90
0.70
0.80
6.0
1.0
0.85
0.80
0.80
0.60
0.60
0.60
0.60
6.5
f
r
5.0 2.5
*
1.9
2.5
*
f
1.0
0.75
* i
0.50
0.50
0.45
0.45
0.45
0.50
it
*
*
*
f f
J = 0.06888
(norhem
0.0003
0.0003
0.0003
0.0006
0.0003
0.0004
0.0003
0.0003
0.0003
0.0010
0.003
0.004
0.015
0.0004
*
*
0.0002
*
mm, +J=OO
0.0006
0.0007
0.0008
0.0004
0.0005
0.0007
0.0006
0.0007
0.0005
0.003 0.0016
*
*
i
*
*
*
*
+
*
f
i
grain size 0.:
Fault Zone, Coastal Cordillera
Ibbro, 30 m N of she ar zone center, 3100
!
4100
70
10 m N of sh ar zone center,
*
nblende,
B7-153
f
i
5510
6589
1180
i
f
820 910
6274
1085
632
5939
1075
315
3186
1070
f
620
2585
1055
f
f
f
f
f
f
undeformed
730
1748
1040
h
2519
10430
1025
2646
965
1000
4253
910
nblende,
789
h
820
87-150
SZI
(*c)
Temp.
Table 4 Argon isotope data for samples from shear zones of the Atacama
10
1.261
1.253
1.272
1.285
1.253
1.309
1.272
1.270
1.229
1.289
1.218
1.196
1.32
1.31
1.10
0.64
1.271
1.267
1.285
1.274
1.250
1.252
1.297
1.25
1.41
1.05
1.32
1.28
1.138
t
0.075
0.070
f
0.027
0.065
f f
0.060
0.070
0.090
0.070
0.060
0.065
0.075
0.080
0.075
0.15
0.20
0.19
0.50
0.020
0.060
0.035
0.040
0.045
0.050
0.060
0.15
0.17
0.40
0.18
0.13
f
_+
*
+
+
*
i
f
*
*
*
*
f
+
f
*
i
f
*
+
It
f
*
*
*
+
40 Ar,d/39Ar,:
17
* f
i *
150
f
*
*
*
f
f
*
*
r
*
f
*
f
f
f
f
*
*
80 3.0
7.5
6.5
8.0
11
8.0
6.0
7.5
9.0
9.5
8.5
17
25
20
60
2.5
7.0
4.0
5.0
5.0
6.0
7.0
20 f
45
20 f
2
151.3
9.0 15
*
f
It
f
152.7
154.1
150.5
156.9
152.6
152.4
147.7
154.7
146.5
143.9
158
157
133
79
152.6
152.1
154.2
152.9
150.1
150.3
155.5
150
168
127
158
153
137.2
Ago
1.6
a.7
4.4
2.6
100.0
99.4
91.8
62.4
77.4
74.6
70.0
23.7
f7.1
12.7
0.5
100.0
86.0
45.0
31.0
15.7
9.8
6.1
5.1
4.2
3.8
3.1
2.0
(%t
C"ill.
j=Ar
21090
47820
25810
22880
22400
21660
5989
11510
855
890
915
940
970
1000
1045
9491
3329
935
970
* f ;
7708
12620
1067
4270
1185
1235
1320
Total
855
590
477.5
299.5
413.5
1055
1110
1190
Total
426.8
5090
4035
815
990
4309
5460
755
775
*
;
*
*
*
*
f
f
*
* f
419.6
1925
715
665
1900
2800
9700
6000
7000
8700
31000
6600
9000
26000
17000
6600
160
-I-
1821
917
3526 3429
3548
4007
7355
3300
4832
6699
8338
4494
*
2.5 40
2.5
0.7
1.3
5
1.5
800
850
950
2100
450
7000
5000
10000
3000
6000
2500
150
1800
6000
2500
1600
4500
2000
f
20.0
12.4
30.0
48.8
21.8
624
779
*
f
*
f
f
*
*
f *
837
f
653
272
26.9
n
0.40
0.05
0.11
0.20
0.09
120
18
140
60
10
0.14
grain size 0.25-0.4
3150
2800
*
f
*
*
*
+
f
*
f
*
*
f
f
80 140
f
*
70
300
450
350 1500
1100
1400
5000
1000
1400
4000
2500
1100
70
2.4
*
7030
9400
5810
7790
9370
6150
3490
9190
2570
7280
2960
1620
5080
300 4000
240
f
*
f i
+
*
*
*
+
f
f
*
+
*
f
*
*
f
*
*
+
i
+
*
*
*
*
f
i
f
*
*
*
+
f
*
*
*
f
*
f
f
f
*
39.8
3.11
58.4
84.3
46.5
100
11s
89
377
77
33.8
f
*
i
i
+
?
*
*
*
*
*
,, J = 0.01395
0.1232
0.1075
0.1081
0.111s
0.1107
0.1626
0.1196
0.1060
0.1137
0.2027
0.2402
0.7789
1.344
4.308 5.063
2.141
1.274
!5 mm,
18.15
2.393
12.23
5.497
11.23
23.57
60.9
28
32.5
37.9
48.3
62.9
46.4
i
16
3.0
19
0.0003
4.0
0.25
3.0
9.0
6.5
60
50
35
210
f
0.0003
0.0003
0.0003
0.0003
0.0004
o.ooo5
0.0002
0.0003
0.0009
0.0016
0.0020
0.004
0.021
0.031
0.007 0.012
f
0.65
0.035
0.35
0.035
0.14
0.90
3.0
12
1.5
2.0
2.5
4.0
2.5
0.090
J = 0.01395
J = 0.06888
mm,
4.326
grain size 0.25-0.5
gain site 0.13-t
75.1 857
700
10000
5500
10000
10000
6000
2550
1000
2500
6500
2000
1500
5000
4000
300
tite. shear zone center
f
f
zt
12350
22230
*
+
*
*
*
*
*
f
*
f
f
1120
87-151
16 440
shear zone CI lter.
*
*
*
*
Y!Z
*
*
*
f
zt
1080
8116
4040
880
1050
1978
835
1498
5558
750
4669
4770
730
1000
561
710
1035
366
660
#7-151
II
*
30720
815
nblende,
42140
775
h
52190
755
TOld
*
27880
715
*
*
459
5254
10 m N of shear zone center,
665
biotite,
585
87-153
1
l-
4409
a4
1189
1352
439.7
130.8
185.8
197.2
308.5
306.3
215.35
314.3
3.2
6.43
31.2
31.8
23.10
106.3
52.1
37.5
6.63
6.07
2.65
2.33
1.53 2.40
0.83
0.29
$689
58.8
231.9
410.9
371.0
368.4
352.0
326.1
510.4
330.5
787.8
579.7
329.5
32.3
k 19
*
*
+
+
f
t
f
*
*
*
f
*
*
+
*
*
f
*
*
*
*
* *
*
f *
*
f
*
f
*
f
*
f
*
It
f
?
*
*
*
2
2.0
2.5
3.0
3.0
3.0
3.0
2.5
3.5
2.5
4.5
3.5
3.0
2.5
30
15
9
9
4.5
3.0
3.0
3.0
3.0
3.0
3.5
6.0
1.1
0.80
1.4
1.5
0.90
4.0
2.5
1.6
0.60
0.60 0.60
0.60
0.60 0.60
0.60
0.60
25
- _-___
* *
1.27 1.26 1.308
f
* *
1.31 1.10
*
6.008
5.898
3.24
6.060
f
i
*
*
f 6.138
f 5.99
f 6.155
6.171
*
*
5.994 6.148
f
4.613
*
f
1.275
1.260
f
*
f
A
1.267
1.286
1.272
1.254
*
*
1.04
1.18
k *
f
1.12 1.04
0.88
*
*
6.160
0.29
f
f
f
f
*
f
+
i
*
*
*
6.21
6.232
6.268
6.361
6.367
6.462
6.320
6.296
6.245
6.223
6.137
f f.
2.18 5.785
0.032
0.60
0.035
0.030
0.055
0.13
0.090
0.085
0.060
0.060
0.070
0.025
0.40
0.17
0.060
0.060
0.050
0.050
0.060
0.055
0.11
0.12
0.30
0.25
0.40 0.25
0.60
0.60
0.020
0.020
0.060
0.035
0.040
0.040
0.055
0.045
0.030
0.045
0.025
0.030
0.045
0.16
*
147.5
138.5
77
142.2
143.9
141.0
140.6
144.3
144.7
144.2
140.7
109.2
151.3
132
157.2
153.0
152.0
154.2
152.7
150.6
156.8
142
152
152
126
125
107 135
36
144.4
145.6
146.0
146.9
148.9
149.1
151.2
f
*
f
f
i
f
f
*
*
*
*
*
f
f
i
f
f
f
*
f
k
*
f
f
f
* *
f
*
f
i
*
*
*
+
i
*
146.3 148.0
f
145.8
*
143.9
t f
52.5 136.0
1.5
14
1.5
1.5
1.8
3.0
2.5
2.5
1.9
1.8
1.9
3.0
45
19
6.5
6.5
6.0
6.0
6.5
6.5
12
14
30
30
30
70 50
30
1.4
5.0
1.9
1.6
1.7
1.7
1.9
1.7
1.5
1.7
1.5
1.5
1.6
3.5
I
6.2
100.0
96.5
70.3
40.6
31.0
28.1
24.1
19.6
13.1
* f f f +
19200
25270
74070
21290
7382
1090
1125
1175
1235
1295
Total
f f
1266 841
Total
)
12
400
350 170
450
400
210
45
160
1
*
1195
550
130
6500
3300
800
f c + * + f f f zt _+ * f +
776
926
2460
7890
11480
11530
25470
19010
24980
26470
73900
13650
11760
730
785
845
905
950
995
1033
1090
1115
1160
1205
1280
Total
f
355
1300
8000
4000
2500
3000
2500
9000
1400
1600
600
120
70
18
30
13.19
4.30
22.32
14.29
10.6
14.72
16.73
66.2
13.37
13.49
16.48
6.85
2.692
2.390
f
f
f
i
*
f
f
i
*
i
f
50.3
808
875 459 _+
* f
+ f
937
*
1094
k
268
f
*
527
274
1.5
300
250 90
300
300
140
18
65
2065
2609
13070
4665
4342
3223
4349
2002
2005
408 1383
230
256
39.2
f
_+
i
+
f
t
f
f
*
zt i
f
f
*
23
15000
7000
400
550
450
1600
250
300
20 110
16
55
2.5
1
I
1 f
f
f
*
it
zt
k
f
ct
0.0003
0.0002
0.0001
0.0001
0.0002
0.0008
0.0026
0.013
/
I
[
200
4.08
5.19
178 4.46
1.62
ze 0.25-0.5
13992
52.0
850.6
760.8
427.7
308.4
200.9
261.1
234.3
233.8
218.8
87.6
24.1
2.1
17.85
1.622
17.59
30.3
12.73
21.44
41.4
17.55
22.15
24.2
11.97
6.296
3.053
1.812
f
i
f
+
*
+
f
i
f
f
f
0.35
0.045
0.12
1.3
0.30
0.40
1.2
0.50
0.90
1.0
0.20
0.090
0.070
f f
0.040
k
I
f
259.4
3992
9.9
451.9
1157.0
620.5
440.4
220.7
311.8
?
*
*
*
f
i
+
k
*
+
*
154.5
*
+ f 324.0
0.40
11
0.55
0.50
11 0.45
22
1.8
3.0
6.0
3.5
2.5
1.8
2.0
1.9
2.0
1.6
1.4
1.3
1.4
J = 0.0139
i
f
i
f f
31.7
7.0
3.2
20
1.9
5.0
4.5
2.5
2.0
1.7
1.9
1.8
1.9
1.8
1.5
1.5
1.4
I 5.569
5.74
5.710
5.745
5.859
5.984
5.762
5.856
5.566
5.514
5.372
4.603
1.516
0.70
f 19 f.
f
f
*
f
f
zt
f
f
f
i
+
f
f
19
5.551
5.12
5.831
5.610
5.684
5.806
5.787
5.612
5.579
5.492
5.306
2.74
1.88
1.51
k
1.12
1.13
1.152
0.83
1.08
1.13
0.42
*
*
+
i
k
i
r
f
*
*
zt
f
f i
*
f
f
i
f
f
f
mm, J = 0.1 6888 f 10
f
i
f
*
*
f
f
zt
+
+
+
k
f
f
.4 mm, J = 0.01395
zone center, grain size 0. 5-t I.7 mm,
0.08908
0.08184 0.08243
0.09295
0.08515
0.3253
0.8564
1.467
0.4623
0.25
0.75
0.25
0.30
1.6
0.35
0.25
2.5
0.50
0.50
0.55
0.12
0.019
* f
0.055
f
)ne center, grain size 0.1:
wdiorite, 5.2 km E ofs ihear zone center,
*
3656
39000 950
*
+
9452
4346
300
200
500
110
130
45
16
8.0
1.1
e, undeformed granor rite, 5.2 km E of she
670
$8-26 bit
f
*
1351
1075
1080
1215
LIC
1556
1045
(
*
1130 1230
f
586
734
1020
T
* f
317
604
910
985
f
1457 +
f
2694
_+
f
1154
2088
i
1198
3227
f
896
+
;
185.3 280
r
29.6
26000
3500
3500
1700
1200
3000
600
700
250
90
30
15
18-26 homlblende, undeformed Q
f
1005
f
+
8719
940 f
*
6904
910
6690
f
4036
875
16080
f
1586
820
12790
*
577
770
1045
f
te, undeformed granod wite. 9 km E of shear
316
bit
710
85-58
sz2
Table 4 (continued)
0.030
0.90
0.035
0.030
0.035
0.035
0.045
0.035
0.040
0.35
0.055
0.12
0.35
0.65
0.11
0.12
0.070
0.20
0.10
0.35
0.25
0.029
0.20
0.035
0.030
0.040
0.040
0.050
0.045
0.045
0.045
0.045
0.080
0.090
0.50
I
I
17
130.6
121
132.5
132.0
133.7
136.4
136.0
132.0
131.3
129.3
125.1
65.7
45.3
36.4
134
136
138.8
101
131
136
52
131.1
135.0
134.3
135.1
140.4 137.6
135.4
137.6
131.0
129.8
128.6
109.0
36.7
i
f
i
f
*
f
k
f
*
f
f
f
f
f
f
f
:
*
f
f
i
f
*
*
* zt
*
f
i
+
f
*
i
f
1.3
2.0
1.4
1.3
1.4
1.4
1.6
1.4
1.4
1.4
1.6
3.0
8.0
15
13
14
8.0
24
12
45
30
1.3
5.0
1.4
1.4
1.5 1.4
1.6
1.5
1.5
1.5
1.5
2.0
2.0
11
1.6 5.2
4.0
100.0
98.0
95.4
93.1
1.0
loo.0
98.6
76.0
55.6
44.7
36.9
31.6
24.8
18.4
12.0
5.6
2.9
0.5
6371 364
4120
1320
1380
Total
I
20460
1120
f
f *
k
f
zt
f
+
*
f
*
f
*
+ k +
6269
8493
15700
28510
19650
820 860
680
890 900
* *
5895
13880
17340
6161
1080
1160
1320
7411
19390
25280
21500
15990
14580
52940
61620
16460
342
940
990
1100
1110
1130
1155
940
975
1005
1055
1280
8441
846
2588
910
528
530
770
Total
zt
+
zt
f
f
f
+_
*
f
f
*
c
=t
fr
*
k
te, mylonite,
840
38-1 e bit
zt
6289 zt
+
9328
*
_+
930 1010
i
f
5443
770
800
* +
f
505
1361 2840
670
730
38-1 c bitHit 8. mylonite.
14200
6355
940
1050
8788
910
6637
54200
890
9365
10740
880
980
2603
820
1020
374
665
850
45
8500
56000
16000
450
10000
9000
5600
2900
700
180
40
25
250
0.9 km
500
4000
1000
280
740
3000
6000
13000
2900
950
2000
1000
160 450
13
1.2 km
400
50
800
4005
4500
1500
1300
1100
1300
40000
630
50
12
I
I
38-1 a bitHit 8, very weakaly defol
*
*
t
*
zk
*
+
*
*
*
+
1.9
130
800
750
250
200
190
200
8200
110
9.0
1.0
f
f
*
f
* *
t
*
+
f
*
*
f
ct f
3.0
86
700
180
130 45
550
1000
2500
500
170
400
180
30 80
1487
8.8
2875
12950
9156
2436
2820
3700
4360
3410
1342
494
274
129.5
66
f
*
k
*
f
+_
zt
*
+
*
+
_+
f
+
*
160
0.90
1500
12000
2900
80
2500
1600
950
500
130
35
12
6.0
26
of shear zone center,
1098
3018
2409
1103 1013
1655
3445
5036
2790
1517
1176
988
258 524
132.7
If shear zone center.
1066
3498
2410
1571
1120
1088
1512
9298
1861
449.9
36.3
ed granodiorite,
i
*
*
f
+
*
f
*
*
*
* *
f * +-
0.12-0.25
+
+
i
*
*
*
*
*
3.71
31.59
0.464
10.3
41.7
105.0
57.5
114.6
31.9
37.8
51.0
227
14.91
6.07
3.557
f
*
*
*
*
f
t
*
*
f
f
*
f
_c
*
0.80
0.020
2.5
2.0
3.0
1.9
0.50
1.5
1.9
2.5
25
0.45
0.15
0.040
0.65
mm,
0.50
0.04
0.16
1.9
0.14
0.20
0.45
350
0.80
0.17
0.27 0.13
0.025 0.43 0.20
mm,
0.14
0.11
0.35
0.30
0.18
0.20
0.14
5.0
0.55
0.30
f c
0.016
r
rain size 0.12-0.25
7.73
11.88
7.93
7.4
5.88
7.75
16.94
539
15.82
8.57
10.58 5.55
1.359 18.92 9.79
ain size
9.01
16.31
13.20
10.43
6.94
6.08
11.04
75.9
13.85
8.95
1.012
I E of shear zone center,
J
.I
L
53.8
158.2
357.2
412.3
246.4
114.9
44.9
20.4
1.8
0.01395
4077
2.1
98.4
207.5
533.8
1825
:
%849
583.0
537.1 832.3
227.6
260.9 116.9
352.4 253.6
330.1
118.8 75.7
71.4
47.2 41.6
0.01395
372.2
386.6
193.4
247.1
146.9
150.9
251.1
264.5
966.3
815.6
14.6
f
f
*
f
*
*
f
f
k
f
*
f
f
zt
f
f
f
f
f *
*
* f
* *
f
f *
*
+ +
19
* 19
k
f
*
*
f
f
It
+
i
*
f
; rai in size 0.12-0.25
22
1.9
1.4
1.7
3.0
10
1.3
1.5
2.5
2.5
1.9
1.5
1.4
1.5
1.3
20
4.5
3.5 4.5
2.0
2.0 1.8
2.5 2.0
2.5
1.8 1.7
1.8
2.0 1.8
3.5
3.0
2.0
2.0
2.0
2.0
2.5
2.5
5.5
5.0
1.9
mm, -r
5.530
5.3
5.622
5.681
5.750
5.864
5.57
5.728
5.728
5.599
5.301
4.641
2.005
1.8
3.5
5.343
5.647
5.640
5.526
5.456 5.434
5.617
5.522 5.603
5.404
5.08
5.209
4.85
1.579 4.13
5.519
5.1
5.698
5.764
5.770
5.771
5.662
5.568
5.615
5.797
5.611
+
zt
*
* It
f
* r
f
*
*
*
* f
*
Lk
*
k
-L
*
k
f
*
f
i
*
It
*
*
f
f
*
*
f
f
f
i
*
f
*
f
f
f
0.01395
5.130
2.17
.I = f
19
0.020
5.0
0.080
0.040
0.025
0.025
0.14
0.050
0.030
0.025
0.035
0.060
0.060
1.3
2.5
0.011
0.035
0.014
0.019
0.080 0.045
0.040
0.030 0.040
0.030
0.11
0.075
0.12
0.070 0.18
0.020
3.0
0.050
0.035
0.060
0.045
0.070
0.070
0.045
0.045
0.020
0.020
0.30 52.1
130.2
125
132.3
133.6
135.2
137.7
131.1
134.7
134.7
131.7
125.0
109.9
48.3
84 43.6
125.9
132.7 132.8
130.1
128.0
132.2 129.5
131.8
130.0
119.9 127.3
122.9
114.7
98.1
38.1
129.9
121
134.0
135.5
135.6
135.6
133.2
131.0
132.1
136.2
132.0
121.1
+
k
f
f
k
*
*
f
*
*
*
*
1.
f
f *
zt
k +
*
2
* *
f
*
f t
f
*
f
r
*
f:
f
*
*
f
f:
f
f
k
f
f
1.3
107
2.1
1.5
1.4
1.4
3.5
1.7
1.4
1.4
1.4
1.7
1.6
59 3.1
1.2
1.3 1.5
1.3
1.5
1.5 2.0
1.5
1.4
3.0 1.4
2.0
3.0
4.0
1.7
1.3
76
1.7
1.5
1.8
1.6
2.0
2.0
1.6
1.6
1.3
1.2
7.0
k +
t t
4005
3491
3272
4227
2564
2063
3935
2119
895
910
935
960
1020
1115
1160
1210 400
600
250
190
1100
350
350
450
650
140
9.5
1.9
f r
2604 3429
f * f
1770 960
t
*
1668
3182
1542
f
i
1973
i
2933
f
3641 3262
f
*
f
1327
362.7
21.1
shs r zone center,
75
500
300
500
190
150
900
300
300
350
500
110
6.0
0.12
gain s
1
1 0.3743
0.0007
1
0.0004
0.1921 r
O.Ogll
0.4377
1
0.0007
0.2569
0.0004
f
35.56
0.005
2.393
0.1361
zt
16.15
0.013
i
* 692.5
+ 59.0
f
* 48.1
116.7
226.8
zt
4.264
30.53
*
k
0.012
34.51
39.26
*
4.723
0.10
33.20
*
f
28.89
i
21.58
10 2.19
0.025
f
8.132
10.12
0.035
0.009
0.005
0.10
f
f
mm, J = 0.068
9.91
6.015
2.967
1.018
t: 0.4-0.7
6.0
2.5
1.4
1.5
4.5
0.45
0.45
0.45
0.45
0.45
0.50
0.50
0.50
0.50
f f
1.137 1.090
f
*
1.093 1.109
f
1.144
f
k
1.146
1.133
f
1.149
+
*
1.147
1.146
*
1.115
i
f
0.731
1.143
+
0.377
0.090
0.010
0.045
0.035
0.015
0.025
0.016
0.030
0.018
0.015
0.014
0.017
0.020
0.017
*
37.7
* *
133.9
f
132.0 136.6
t
* 137.8
138.1
*
*
31.5
1 38.5
*
1 37.1
f
f
38.2
1 38.1
*
134.4
i mt
46 89.3
1.3
5.5
4.0
1.8
3.0
1.9
4.0
2.0
1.8
1.7
2.0
2.5
2.0
11
100.0
91.7
64.6
66.3
36.5
31.6
29.3
25.1
20.0
14.4
9.6
5.7
0.9
Analytical methods: Irradiation (hornblendes: 5 days, biotites: 24 hours) of the samples was carried out under Cd shielding at the Kemforschungszentrum Geesthacht, Gas extraction and purification followed the procedure of Hammerschmidt et al. (1992). The reciprocal sensitivity of the mass spectrometer was determined as 1.85 *0.005X lo- “I cm’/mV using an Ar standard of known volume. The extraction blank for “‘Ar was 2.50f0.25X IO-’ cm’ STP between 400” and 1100°C and increased linearly to 1.5X 10e8 cm’ STP between 1100” and 1500°C. For masses 36, 37 and 38 the measured blanks were = 2 X IO- ” cm? STP, while for mass 39 it was = 3X IO-” cm’ STP “Ar equivalents. The errors in Table 3 are at the 95% confidence level and include components from the statistics of measurements, discdmina6on, blank (usually 20% for j9Ar and 30% for other isotopes), interferences, reciprocal sensitivity, weighing errors and relative measurements of the neutron flux.The following interference corrections were used:“hAr/ “Ar(Ca) = (2.7 + 0.2) X 10-J3XAr/ “Ar(Ca) = (6.0 5 2.0) X 1O-““Ar/ “Ar(Ca) = (6.8 f 0.2) X lo- j4”‘Ar/ “AI = (3.0+3.0)X 10-4i8Ar/‘9Ar(K) = (1.4*0.3) X IO- ““Ar/ ‘“Ar(K) = (6.0 7t 2.0)X IO- “‘Ar(K) = neutron-induced only from K1.U.G.S. constants (Steiger and Ilger, 1977)LP-6 biotite standard (Engels and Ingamells, 1971)
It
k
k
I
A
4471
875
f
1775
f
560.7
800
855
f
303.5
lblende, mylonite,
730
Bf3-lg hc
Table 4 (continued)
E. Scheuber et al./Tectonophysics
250 (1995) 6/-87
77
r 0740 3j E 0 720
87-151 WR plg-bi
87-151 WR-pIg-hb
MSWD
= 148.5 f 7.9 Ma
= 143.9 i 0.3 Ma
= 0.703466 * 32 = 0.22
= 0.703490 f 30
. . : . . . : . ,
0.7030.
0.2
0
: .
4
06
0.4
0
40
20
60
60
Fig. 6. Rb-Sr isochron diagrams of minerals and whole rocks from the shear zone SZl; the insert in the upper left corner of each diagram is an enlarged view of the lower left comer. M refer to data included for age calculation, 0 data not included for age calculation. WR = whole rock; pig = plagioclase: LM = light minerals (plagioclase plus some quartz and K-feldspar); hb = hornblende; bi = biotite.
Table 5 Compilation of all age results (for “Ar/ these steps is indicated)
T
Sample
j9Ar plateau ages the number of consecutive
Number
Fiatmu
Total
of steps
“Ar
A-
releesa
heating steps and the amount of ““Ar released during
Rb-Sr (Ma) mean age (-Ma)
I%1
SZl
_..
87-150
hb
152.6
f
2.4
152.9
f
2.4
12
98.0
87-153
hb
151.3
zk 3.2
151.7
i
3.2
15
99.5
87-l 53 bi
144.4
f
1.4
149.3
f
1.4
5
29.2
150.3
87-151
hb
151.3
f
2.9
151.7
i
2.8
16
99.6
148.5
2 7.9
87-151
bi
138.5
f
1.5
142.8
f
1.4
8
90.3
143.9
f
0.3
44.2
129.1
f _--
0.4 0.3
* 0.3
-i_
s22 85-58
bi
131.1
f
1.3
134.8
f
88-26
hb
137.1
f
7.2
138.6
? 75
88-26
bi
130.6
k
1.3
132.7
55
129.3
f
88-1s
bi
129.9
1.3
134.8
33.9
129.0
* 0.3
hb
f ___
1.0
88-1~
f --_
133.6
f
1.2
88-1~
bi
125.9
+
1.2
131.9
*
1.2
48.3
125.3
f
0.3
88-10
bi
130.2
+ 1.3
133.7
f
1.2
88-10
hb
133.9
*
138.0
k 1.7
1.3
1.3
94.8
zk 0.7
19.9
)_
74.9
129.4 126.1
t 0.3 + 0.6’ * hc
ksnde and biotite
78
E. Scheuber et al./ Tectonophysics 250 (1995) 61-87
obtained by laser spot age determinations indicate a homogenization effect because during vacuum degassing by conventional heating the lattice of hydroxyl-bearing minerals like hornblende (Lee et al., 1991) and biotite (Hodges et al., 1994) breaks down to form oxy-hornblende and oxy-biotite and at higher temperatures metal oxides. These reaction products display an @Ar/ 3gAr distribution different from the original minerals. (4) But also “Ar recoil due to the neutron reaction with K may be responsible for staircase age spectra in two ways. Firstly, “Ar is lost during irradiation from the mineral and as a consequence the age seems to be too old (Hess and Lippolt, 1986). Secondly, “Ar recoil from only slightly altered biotite (Hess et al., 1987) to K-poor mineral phases like chlorite or Fe-Ti oxides and titanite which sometimes can only be identified by TEM methods leads to age spectra whose inferred ages are geologically too high and thus not meaningful even if well-defined plateau ages are given. Thus, only independent isotope age determinations corroborate the reliability of the ages. In this respect it is plausible to test our “OAr/ 3gAr results by Rb-Sr determinations whenever it was possible. 4.1. Age results of the amphibolite facies shear zone (SZI J 4.1 .I. Hornblende
From the undeformed gabbro through the margin to the center of the shear zone hornblende yielded remarkably constant plateau ages of 152.9 + 2.4 Ma (sample 87-150) 151.7 f 3.2 Ma (87-153), and 15 1.7 f 2.9 Ma (87- 15 l), respectively (Fig. 8). These ages agree with the total degassing ages (Tables 4 and 5) and, together with the amount of “Ar released (98-99.6%, Table 5), they indicate that no significant Ar-loss has occurred since the closure of the system. Rb-Sr data on hornblende from the shear zone center (87-151) gave an age of 148.5 + 7.9 Ma (Table 3), which is in agreement with the 40Ar/ 39Ar ages. However, from the undeformed and moderately deformed samples (87-150, 87-153) no Rb-Sr homblende age information can be obtained. 4.1.2. Biotite The 4oAr/ 39Ar as well as Rb-Sr ages of biotite display a pattern completely different from the hom-
blendes. Biotite from the weakly deformed sample 10 m north of the shear zone center (87-153) gave a Rb-Sr age of 150.3 _t 0.3 Ma which agrees with the “OAr/ 39Ar plateau age of 149.3 + 1.4 Ma (Fig. 8). But the lower total degassing age of 144.4 + 1.4 Ma together with the staircase pattern of this age spectrum reveal an Ar-loss of 5.3%. From the shear zone center (87-151) still younger age values have been determined on strongly recrystallized biotites which yielded a Rb-Sr age of 143.9 f 0.3 Ma (Fig. 6) and a plateau age of 142.8 + 1.4 Ma (Fig. 8). In this sample also an Ar loss is indicated by a younger total degassing age of 138.5 -+_1.5 Ma. The finding that the Rb-Sr biotite ages are equal to the plateau ages but older than the total degassing ages suggests thermal overprinting which can also be inferred from the statically recrystallized fabrics of the samples (Section 3). Radiogenic Ar is lost more easily than Sr, and this Ar loss can be detected only by the *Ar/ 3gAr method. The staircase age pattern is not a result of very slow cooling (< l”C/Ma), because in a sample, collected close to SZl, fission track age dating on apatite (Scheuber and Andriessen, 1990) indicates an age of 110 Ma. Taking the closure temperature of apatite (100°C) and of biotite (300°C) and their age values into account an integrated cooling rate of S.O”C/Ma is inferred. This value is approximately one order of magnitude higher than that of a very slowly cooling enviroment. Obviously, vacuum degassing and recoil does not affect the 40Ar and 3gAr distribution since the plateau ages are in accordance with the Rb-Sr ages. 4.2. Age results from
the greenschist
facies
shear
zone (SZ2) 4.2.1. Hornblende
Euhedral hornblende from the undeformed granodiorite gave a total degassing age of 137.1 f 7.2 Ma (sample 88-26, Fig. 9). From the age spectrum the highest obtained age value amounts to I38 + 7.5 Ma. Spotty homblendes from the shear zone center confirm this age: the age spectrum of hornblende (88-lg) reveals an age of 138.0 + 1.7 Ma (Fig. 9). But the staircase pattern also indicates an Ar loss of 5.3%. Therefore the total degassing age is lower at 133.9 + 1.3 Ma. The Rb-Sr data of the undeformed rocks (Fig. 7) show a high degree of isotopic disequi-
E. Scheuber et al./ Tectonophysics 250 (1995) 61-87
bi _,,=.-
,
0
80
60
40
20
0
20
40
60
loo
80
120
$10
“Rb?%r
“Rb/%
1
m-.-_r”.“-
0.711
0
20
40
80
60
100
0
1
4
3
2
“Rbl%r
“RbPCN
0.820
0.760
I
t
6 0.780 % 3j 0 0.740
02
04
0.6
_'
= 129.4 L 0.3 M8 = 0.70344a f 4s MSWD 10.24
0.700 0
20
SO
40
100
60
*‘RW”Sr
o
I
2
3
4
0
10
20
30
40
50
8’RbPSr
5
6
7
“Rbl%r Fig. 7. Rb-Sr isochron diagrams
of minerals and whole rocks from the shear zone SZ2; for explanation
see Fig. 6.
60
80
E. Scheuber et al./
Tectonophysics 250 (1995) 61-87
librium between hornblende and the other constituents. In the 87Sr/ 86Sr vs. 87Rb/ 86Sr diagram (Fig. 7) and in the reduced isochron plot (Fig. IO) hornblende shows isotope disequilibrium; hence no Rb-Sr age information could be gained. By contrast, in the mylonite sample (88-1~) whole rock, the light mineral fraction and hornblende define an isochron whose slope corresponds to an age of 133.6 + 1.2 Ma with 87Sr/ “Sr initial ratio of 0.703398 + 20 (MSWD = 0.07). 4.2.2. Biotite The undeformed granodiorite (sam les 85-58 and B 88-26) yielded identical Rb-Sr and Ar/ 39Ar total degassing ages between 13 1.l and 129.1 Ma (Table 5; Figs. 7 and 9). From the weakly deformed gran-
152.9 f 2.4 Ma
odiorite to the shear zone the biotite ages drop from 129.9-129.0 Ma (88-la) to 125.9-125.3 Ma (88-1~). The same age (126.1 + 0.6 Ma, Sri = 0.703515 + 40) results from an isochron defined by whole-rock, light minerals and hornblende containing - 30 wt.% small (< 20 pm> biotite flakes (88-lg). Within the shear zone sample 88-le still preserves the higher age of 130.2 + 1.3 Ma cmAr/ 39Ar> and 129.4 + 0.3 Ma (Rb-Sr), respectively. Comparing the 4oAr/ 39Ar and Rb-Sr age results from SZl and SZ2 we find in SZ2 a sequence of age values completely different from those of SZl. In SZ2 @Ar/ 39Ar plateau ages are generally higher than Rb-Sr ages, whereas the total degassing ages agree perfectly with the Rb-Sr ages (Table 5). According to Roddick et al. (1980) and Foland (1983)
87-150 hb >
total degassing age: 152.6 i 2.4 Ma 20
40
60
80
100
per cent of “Ar released
3 180-z.
67-153 biotiie
87-l 53 hb
P $ H m
140-*-
151.7 f 3.2 Ma 5’
100.
0
>
I’y
:
total &gassing age: 151.3 ) 3.2 Ma 20
40
60
80
100
0
20
per cent of “Ar released
:
;
:
total degassing age: 144.4 f 1.4 Ma
40
80
80
j 100
per cent of “Ar released
7 z
180
0 E E
140 total degassing age: 138.5 f 1.5 Ma
H lQ loo 0
20
40
60
80
100
per cent of “Ar released
Fig. 8. “Ar/
39Ar age spectra of samples from the amphitdite
20
40
60
per cent of “Ar released
facies shear zone (SZl).
80
100
E. Scheuber et al. / Tectonophysics
250 (I 995) 61-87
81
180 c 160 5 140 8l n 120 80-r
I!
60 -I* 40 + 0
total degassing age: 137.1 f 7.5 Ma
E
100
5a 0
8o 60
40
60
80
total degassing age: 133.9 f 1.7 Ma
40 4 0
1 20
;j
100
I 20
per cent of “‘Ar released
40
60
80
100
per cent of 3eAr released
88-26 biotite
132.7 f 0.7Ma
total degassing age: 130.6 f 1.3 Ma
total degassing age: 131.1 t 1.3 Iha 20
40
60
80
20
100
per cent of “Ar released
40
60
80
100
per cent of “Ar released
5
% 80 60--
‘i assing age: 129.9 f 1.3 Ma 40
60
80
n 100
0
per cent of “Ar released
60
total degassing age: 125.9 f 1.2 Ma
40+ 20
40
60
80
100
per cent of “Ar released
total degassing age: 130.2 f 1.3 Ma
20
40
60
80
100
per cent of “Ar released
Fig. 9. “OAr/ j9Ar age spectra of samples from the greenschist
elevated plateau ages were related to excess argon; however, in their cases the total degassing ages of micas were higher than the Rb-Sr ages due to the incorporation of excess argon. By contrast, in our samples the concordance of total degassing and Rb-Sr
facies shear zone (SZ2).
ages excludes both explanations: excess argon or loss of 39Ar during neutron irradiation. Although our biotite samples were microscopically homogeneous the only explanation for the Rb-Sr and “OAr/ 39Ar age discrepancy we can think of is that given by
82
E. Scheuber et al./
Tecionophysics 250 (1995161-87
55
87-151
"0 '- 35 g VI r x 15 .5
‘g
WR
0.01
0.1
bi
$b
il. 0.01
0.1
1
10
100
%b/-Sr
0.1
1
10
1
10
100
10
100
*'Rb/%r
100
1000
0.01
0.1
1 B'Rb/%ir
87Rb/%r
Fig. 10. Deviation from the isochrons (solid lines) of samples from SZl (87-153 moderately deformed, 87-151 strongly deformed) and from SZ2 (88-26 undeformed, 88-lg S-C mylonite and 85-54b ultramylonite). Deviation = [(s’Sr/ *6Sr,)/(87Sr/ s6Sr,, - l}] x lo5 where the suffix m and BF denotes the measured and the best fit values, respectively. WR = whole rock; bi = biotite; hb = hornblende; plag = plagioclase; LM = light minerals.
Hess et al. (1987).These authors have demonstrated that microscopically perfect biotites with only a small K-deficit due to altered inclusion phases like chlorite and Fe-Ti oxide s, only visible with TEM methods, reveal staircase age patterns and plateau ages due to recoil 39Ar which is redistributed from K-rich lattice sites to K-poor ones. The plateau ages are higher than the geological constraint ages and a K-content of 8% and more seems to be necessary to circumvent such problems. As a further consequence of the ‘“Ar redistribution we find sometimes very low age values in the first degassing steps which, however, do not indicate Ar loss by thermal overprinting, because, from the mass balance point of view, it is obvious that ,me plateau-like age values which are higher than the Rb-Sr ages have to be compensated by low apparent ages in the low-temperature degassing steps. From the observation that the K-contents of the biotites from SZ2 were always smaller than 7.75
wt.% (microprobe analyses) and that the Rb-Sr ages was lower than the plateau age we are probably faced with the problem of 39Ar redistribution. Therefore, in the case of biotites from SZ2 we consider total degassing ages only. 4.3. Redistribution
of Sr isotopes
The diagrams of Fig. 10 show the deviation in parts per lo5 of Rb-Sr data from the isochrons. In SZl from the moderately deformed sample (87-153) plagioclase and hornblende show a deviation from the whole-rock-biotite reference line of some 110 and 500 ppm, respectively, hence these data were omitted. By contrast, the data of sample 87-151 (shear zone center) only show deviations of 3-39 ppm indicating a very strong isotopic homogeneity of the rock. In SZ2 a similar homogenization of the Sr isotopes can be observed. In the undeformed sample (88-26) hornblende shows a deviation of
E. Scheuber et al. / Tectonophysics
- 71 ppm from the isochron; by contrast, in the mylonitic sample (88-lg) the deviation of homblende + biotite is reduced to - 1 ppm. Furthermore, from an ultramylonitic band from the shear zone center (sample 8554b) we obtained a very low 87Rb/ 86Sr ratio of 0.02 (Table 1); this agrees with the observed lack of any K-feldspar, biotite and hornblende. The measured 87Sr/ 86Sr ratio amounts to 0.703551 &-70. This sample fits the isochron of sample 88-lg perfectly, indicated by a deviation of only - 2 ppm. In particular a recalculation of the isochron (88-lg hb + bi, LM, WR) including 85-54b whole rock gave an isochron age of 126.1 f 0.6 Ma (MSWD = 0.39) with an initial 87Sr/86Sr ratio of 0.703515 f 18.
5. Discussion: geological interpretation of the ages 5.1. Atnphibolite facies shear zone (SZI 1 The high degree of consistency of the age data reported in the previous chapters is illustrated by the fact that 4oAr/ 39Ar and Rb-Sr methods gave identical ages and that each age range has been obtained from at least two samples; e.g., all biotites of undeformed samples in SZ2 yielded identical 40Ar/ 39Ar and Rb-Sr ages. Uniform *Ar/ 39Ar plateau ages of 153- 152 Ma (Table 5) have been obtained from igneous as well as from completely recrystallized homblendes irrespective of the state of strain across the shear zone. Thus, we interpret the hornblende age as a cooling age due to the regional cooling of the gabbro and its surrounding rocks. Hornblende dates the time when the shear zone together with its host rock passed the isotherm of the hornblende closure temperature; on the other hand, the age of the first deformational event which produced the recrystallized homblendeplagioclase fabric is postdated. Furthermore, the biotites from the moderately deformed sample are about 2-3 Ma younger (149.8 Ma, geological mean age) than hornblende (Table 5); this is in accordance with the concept of closure temperatures where homblende (T, _ 530°C Harrison, 1981) closes prior to biotite (T, - 300°C Purdy and Jager, 1976; Harrison, 1985; van Blanckenburg et al., 1989; Hurford et al., 1989). From the age difference between hornblende
250 (1995) 61-87
83
and biotite the rate of cooling can be calculated as 70- 11 S”C/Ma. In the shear zone center (87-15 1) the Rb-Sr and the @Ar/ 39Ar ages of biotite are completely reset due to the recrystallization of this mineral. The ages amount to 143.9 &-0.3 Ma (Rb-Sr) and 142.8 + 1.4 Ma (“OAr/ 39Ar). This complete age resetting and the homogenization of the Rb-Sr system indicates that the time of deformation 143.4 + 0.3 Ma ago is determined and that it is not a mixed age. From the fact that biotite was reset at 143.4 Ma whereas homblende shows no resetting, it can be concluded that this second deformation was not strong enough or that the temperatures were not sufficient to affect hornblende. The Ar-loss of 3.0% in sample 87-15 1 is interpreted as being the result of static recovery of the microstructure which, however, was not sufficient to disturb the Rb-Sr system. It is important that this recovery affected only the shear zone center and not the moderately deformed sample; thus, static recovery was rather due to a relatively slow stress drop after deformation than to later heating, as in the latter case static recovery should also have affected the moderately deformed sample 87-153. In summary, from the age relations in SZl the following conclusions on the course of crystallization and deformational events can be drawn: (1) Before 152- 150 Ma (a) the intrusion of the gabbroic protolith took place, followed (b) by the development of the recrystallized plagioclase-hornblende fabric due to sinistral strike-slip movements, and (c) by the static growth of biotite (only detectable in the moderately deformed sample (87-153). (2) At N 143 Ma the second deformation (dip-slip normal movements) occurred which led to the recrystallization of biotite in the shear zone center which (3) was also subject to minor static recovery after 143 Ma. 5.2. Greenschist facies shear zone (SZ2.J The oldest age value of SZ2 (139- 138 Ma) is yielded by the @Ar/ 39Ar spectra of hornblende from the undeformed granodiorite as well as from homblende porphyroclasts from the interior of the shear zone. This is interpreted as the cooling age of the igneous rock below the closure temperature of homblende. Rb-Sr hornblende ages show partly negative or geologically unreasonable age values; they do not
84
E. Scheuber et al. / Tectonophysics
give an age information because they are not in isotopic equilibrium with other rock constituents. The biotite ages of - 130 Ma (Rb-Sr and “OAr/ 39Ar) are also interpreted as cooling ages of the undeformed granodiorite. A cooling rate of - 2030”C/Ma can be calculated from the age difference between biotite and hornblende. The younger biotite ages of 126- 125 Ma (Rb-Sr and 4oAr/ 39Ar) are restricted to the mylonites; this suggests that this age is due to resetting during mylonitization. From the fact that resetting was associated with a homogenization of the Sr isotope system it is concluded that resetting was complete in the samples which gave 126- 12.5 Ma. Thus, the reset age is the deformation age and not a mixed value. Although resetting was complete in single samples not all mylonites show the same degree of resetting. Sample 88-le shows the original cooling age of biotite ( - 129.8 Ma); this is interpreted as a relic age (see below). The age relations in SZ2 can be summarized as follows: (1) cooling of the igneous protolith below c of hornblende occurred at - 138 Ma; (2) cooling below c of biotite at - 130 Ma; (3) shear deformation took place at 126- 125 Ma, and no static recovery after deformation can be observed. 5.3. Age resetting by deformation According to the criteria outlined in Section 1 in each shear zone the age of one deformational event could be determined because (1) the same mineral species is younger in the deformed than in the undeformed rocks, and (2) the recrystallized minerals are in isotopic equilibrium. In SZl the first deformation took place above the closure temperature of the dated minerals; this deformational age is therefore obscured by the regional cooling. Consequently only a minimum age for this deformation could be determined. On the other hand, the younger deformation (at - 143 Ma) could be dated as it is restricted to the deformed rocks. Deformation thus started below the closure temperature of the host rock. The homogenization of Sr shows that resetting was complete and thus the obtained age is not a mixing age. Further resetting after deformation due to partial static recovery had only the minor effect of some loss of At-. In SZ2 deformation also started below T,
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of the protolith and age resetting was also complete, indicated by Sr homogenization in the shear zone center. Furthermore, in this shear zone no postdeformational loss of Ar occurred due to the rapid stress drop after deformation. The patterns of age resetting and homogenization of Sr isotopes can be interpreted in terms of the influence of finite strain, strain rate, stress, deformation mechanisms and temperature: resetting and homogenization is most complete in the most strongly deformed samples from the centers of the shear zones. This suggests an influence of the finite strain on the resetting behaviour. Weakly or moderately deformed samples may show relic ages or only partial resetting. However, not only the finite strain but also strain rate and stress and probably deformation mechanisms have influenced the age resetting. One observation which indicates that stress and strain rate influenced age resetting is the relic age of the mylonitic sample 8%le which has a lower quartz content than the other samples except for the ultramylonite 8%54b (Table 1). The low quartz content resulted in a higher strength during deformation, and, consequently, a lower strain rate. Another observation illustrating the influence of deformation mechanisms is the age resetting and complete Sr-isotope homogenization in the ultramylonite sample 85-54b. The very fine grain size and the compositional changes that affected this rock suggest that diffusional mass transfer was operative as one deformation mechanism, which is a very probable reason for isotopic homogenization. Although age resetting and homogenization can be explained by deformation alone, a further influence of shear heating cannot be excluded. In SZl shear heating can be concluded from the presence of metamorphic orthopyroxene which is limited to the center of SZl. In SZ2 the increasing degree of dynamic recrystallization of feldspar in the more strongly deformed samples may be due to an increase of temperature.
6. Geological implications We have obtained tight age constraints for the timing of movements along the Atacama Fault Zone. In SZl the cooling age of homblendes ( - 152 Ma) -which is very common as an intrusion age in the
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Coastal Cordillera (Maksaev et al., 1988; Her& and Marinovic, 1989; Boric et al., 1990)--corresponds to the Araucanian tectonic phase which took place in the Kimmeridgian (Riccardi, 1988), and which resulted in a > 60” angular unconformity between the Jurassic arc lavas and very coarse overlying conglomerates of Kimmeridgian to Early Cretaceous age. On a regional scale the Kimmeridgian deformations in the magmatic arc are contemporaneous with a change from marine to continental sedimentation in the eastward adjacent backarc area @rinz et al., 1994). The age resetting at N 143 Ma has more local importance; it corresponds to vertical movements around the Pluton Cristales. The deformation age of SZ2 gives the age of the strongest sinistral movements along the AFZ. In contrast to the Late Jurassic shear zones, the occurrence of Early Cretaceous mylonites is longitudinally much more constant. The ages of SZ2 coincide with several fission track data obtained from the Coastal Cordillera (130- 100 Ma; Maksaev et al., 1988; Maksaev, 1990; Scheuber and Andriessen, 1990). The deformation age in SZ2 thus coincides with the onset of cooling and uplift of the Jurassic-Early Cretaceous magmatic arc. Uplift and cooling were followed, at about 110 Ma, by the eastward shift of the arc into the Longitudinal Valley (Scheuber et al., 1994). Acknowledgements This work was part of the project “Mobility of Active Continental Margins” and also of the Sonderforschungsbereich 267 (“Deformation Processes in the Andes”) supported by the Deutsche Forschungsgemeinschaft. Thanks to Prof. Reutter, Berlin, and to the colleagues from the Universidad Catolica de1 Norte, Antofagasta for joint field work and discussions. We express gratitude to the staff of the Kemforschungszentrum Geesthacht for facilitating fast neutron activation. We are indebted to Dr. S. Teufel and M. Feth for supporting the experimental work. References Boric, R., Diaz, F. and Maksaev, V., 1990. Geologia y yacimientos metaliferos de la Regi6n de Antofagasta. Serv. Nat. Geol. Min. Bol. 40, 246 pp.
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