Brittle Fault Activity Constrained by K-Ar Ages

In this study, we used the different illite polytypes to decipher different illite-forming events, thus, the time span of the deformation history. In diagenetic to very low grade sedimentary rocks, 1Md and 1M polytypes are considered authigenic products formed under diagenetic to anchimetamorphic, prograde conditions during subsequent burial (e.g. Grathoff and Moore 1996). In contrast, the 2M1 illite polytype is considered a detrital component due to its restriction to epizonal conditions.

However, at higher temperatures, even 2M1 illite could be developed in fault gouges, especially if the onset of brittle deformation and, thus, fault gouge development directly follows cooling to brittle deformation temperatures (about 300°C for quartz, e.g. Passchier and Trouw 2005 and references therein). The development of 2M1 illite polytypes in a brittle fault gouge is possible due to cooling and its passage through epizonal conditions during retrograde metamorphism of the host rock, contemporaneous to faulting (Fig. 4.4). Thus, in contrast to a sedimentary environment, the 2M1 illite must not be excluded from consideration but can be considered to record the onset of fault gouge formation.

Based on the calculated polytype compositions of the samples, we extrapolated the ‘end-member’

age of the 1Md polytype and the 2M1 polytype (hypothetically samples which consist of 0% 2M1 illite and 100 % 2M1 illite, respectively) by plotting the age of each individual grain-size fraction of a fault gouge sample against the 2M1 illite content (Fig. 4.9 and Table 4.1). These plots show a coefficient of determination (R²) always larger than 0.9, confirming a clear linear relationship between age and 2M1 polytype content. In this study, the 1Md end-member age can be interpreted to best represent the age of the youngest movement because the age-increasing influence of the “older” 2M1

polytypes is eliminated, whereas the 2M1 end-member age representing a) the oldest generation of neoformed illite and/or b) the age of ”detrital” muscovite, meaning crushed muscovite from the host rock. Contamination of mineral fine fractions (<2 µm, <0.2 µm) by cataclastically crushed muscovite from the host rock is very unlikely because of the very strong mechanical resistance of this mineral.

Muscovite flakes are apt to rotate parallel to the faulting plane rather than being ground into extremely small particles (e.g. Wemmer 1991). If so, they could be identified by their excellent crystallinity (ca. 0.060 Δ°2θ).

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Fig. 4.9: K-Ar ages vs. % 2M1 illite content for three analyzed samples.

Errors on polytype age determinations are largely dependent on the reliability of illite compositional analysis. Grathoff and Moore (1996) conservatively assume an error of 2-5% for the polytype quantification method. However, they performed polytype quantification on nearly pure illite samples (see Grathoff and Moore 1996). In this study, one of the main problems in polytype quantification was the abundance of kaolinite, as well as substantial amounts of smectite. Especially XRD reflections from smectite exhibit overlap with illite polytype specific reflections. As a consequence, the illite content might be overestimated in the case of superposition of smectite reflections on illite peaks. Polytype quantification obtained by this study could be quantified (Table 4.3) but is subjected to errors greater than those reported by Grathoff and Moore (1996).

Thus, polytype quantification results can only be considered as rough estimations due to the abundance of interfering phases. As a consequence, extrapolated 100 % 2M1 and 0 % 2M1 illite ages are subjected to substantial error sources. Extrapolated ages are reported here (Table 4.1) but are excluded from interpretation. Nevertheless, polytype quantification is in very good accordance with other parameters, such as grain-size age, illite crystallinity and K-Ar age, indicating the consistency of the data set. Even most of the extrapolated polytype end-member ages (Table 4.3) are in good accordance with K-Ar mica cooling ages (see below).

Following the above-stated assumptions, we interpret all illite to be neoformed, i.e. to be fault-gouge related. However, non-deformational illite formation by fluid percolation cannot be excluded but is unlikely due to the consistency of the data set. The wide age span of the dated sample fractions documents a long-lasting fault activity from 315 Ma (Early Carboniferous) to 179 Ma (Early/Middle Jurassic), whereby the relationship of increasing K-Ar ages with increasing grain size (Table 4.1 and Fig. 4.5) is consistent with increasing content of the 2M1 illite polytype, formed in the earlier fault history under higher temperatures. However, especially the larger grain-size fractions have to be considered as mixtures of illite formed at different times during different events and, thus, to show an age younger than the effectively oldest illite-forming event. Additionally, the youngest age

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documented by fault gouge dating cannot be considered as the cessation of fault activity but to represent the last illite-forming event and, thus, the cooling below illite-forming temperatures (see Fig. 4.3 and discussion below).

When the 2M1 illite age represents the onset of brittle deformation, illite ages should always be younger than K-Ar biotite cooling ages (Fig. 4.4). Biotite shows a closure temperature around 350°C (Harrison et al. 2009) and can be interpreted to date the cooling of the basement from ductile to brittle deformation temperatures. Thus, biotite ages represent cooling ages before the first brittle deformation and formation of illite containing fault gouges (Fig. 4.4).

In the Nogolí region, K-Ar biotite ages document a basement cooling to epimetamorphic temperature conditions during the Middle Mississippian (Varela et al. 1994; Sims et al. 1998;

González and Sato 2000; Sosa et al. 2002; Steenken et al.2004; 2008). Further constraint is given by K-Ar and Ar-Ar muscovite ages (Sims et al. 1998; Steenken et al. 2008) taken from several mylonitic shear zones in this region, yielding ages of 380-350 Ma. These ages are interpreted as the last mylonitization event caused by the Achalian Orogenic Cycle before the final cooling of the basement to brittle deformation temperatures (below 300°C for quartz, e.g. Passchier and Trouw 2005 and references therein; see also Fig. 4.4).

In addition, the ages obtained by this study are consistent with epimetamorphic illites from the San Luis Formation (330-290 Ma; Wemmer et al. 2011). The latter are associated with the Toco orogeny (Bahlburg and Breitkreuz 1991).

Whether our illite fine-fraction ages are influenced by this event or not is unclear. We think that Late Carboniferous ages from both locations can be better related to the last episode of the Chanic phase of the Famatina Orogeny, which reactivated several shear zones in the Sierras Pampeanas (e.g.

Martino 2003).

Illite-generating fault gouge activity along the sampled faults is interpreted to have ceased between 222 Ma (APG 50-09 <0.2 µm) and 173 Ma (APG 59-09 100% 1Md), as shown by the majority of the

<0.2 µm grain-size fractions as well as the calculated 100% 1Md illite fractions. Only sample APG 6009 shows a younger age for the 100% 1Md fraction of 119 Ma. Similar to the unusually high age of the calculated 100% 2M1 fraction of sample 59-09 discussed above, this comparatively young age might be related to an extrapolation error resulting from its – compared to the other samples – low 1Md illite content in the <0.2 µm fraction (see Table 4.1). The youngest age documented by fault gouge dating must not be considered to represent the cessation of fault activity but to represent the last illite-forming event and, thus, the cooling of the fault block below illite-forming temperatures (approximately between 75-110°C; e.g. Hamilton et al. 1992; see Fig. 4.4). Cooling below the illite-forming temperatures is constrained by AFT and AHe ages (Tables 4.4 and 4.5). The youngest illite must overlap with the apatite fission track ages (representing cooling below 110°C), whereas the AHe

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ages (representing cooling below 60°C) must always be younger than the K-Ar illite ages (Fig. 4.4).

This can indeed be observed for all analysed samples (Tables 4.4 and 4.5, Figs. 4.5 and 4.6) 4.6.2 Thermal Modeling

Fission track single grain ages and confined track length distributions are used along with the (U-Th)/He ages to numerically model possible t-T paths for individual samples. Eight samples along the transect were modelled following the approach of Ketcham (2005) using HeFTy 1.7.0 software.

Several boundary conditions were set to the thermal models: (1) the age and temperature range of the K-Ar ages from brittle fault gouges constrains the beginning of transition from ductile to brittle temperatures and, thus, to <300°C due to the onset of brittle behaviour of quartz below this temperature (e.g. van Daalen et al. 1999; Passchier and Trouw 2005 and references therein); (2) the age of the pure 1Md illite fraction (Table 4.3) from fault gouge dating indicates cooling below the illite-forming temperatures (approx. 75110°C), so reasonable models should show cooling below this temperature range before the youngest 1Md illite-forming event; (3) the annual mean surface temperature of 17 °C (Müller 1996) defines the end of the time-temperature path. The resulting cooling models are shown in Fig. 4.10.

Based on the individual cooling paths derived from HeFTy modelling (Fig. 4.10), a regional thermal history for the entire transect was compiled and is shown in Figure 4.11. This compilation highlights a similar thermal history for all samples within the transect. Within a margin of error, ages are similar for the passage of different temperature regimes for individual samples and respective dating systems. Based on the upper temperature boundary defined by the ZHe ages, cooling below the PRZZ

temperatures (≈175 °C) over the whole transect started in Latest Carboniferous to Middle Permian times. An exception of this is given by samples 49-08 and 34-08, which show initial cooling below the PRZZ in Carboniferous times. This clear deviation from the trend shown by all other samples gives an indication of an older thermal history of these samples compared to the others. The temperature regime for the apatite fission track partial-annealing zone (PAZA ≈110-90 °C) was passed in Middle Permian to Early Triassic times. The lower temperature boundary recorded in our data (PRZA ≈65 °C) was reached in Late Permian to Jurassic times, and, for some samples, even Cretaceous times (Fig. 4.11). The span of this time interval is related to very slow cooling during this time, as is evident in all individual models (see Fig. 4.10). In combination with a long-lasting passage through the PRZZ, PAZA and PRZA, this led to a broad scattering of helium and fission-track ages (see Fig. 4.7). The long-lasting passage through the PAZA is also evident in the broad, unimodal, fission track length distribution and distinct, shortened tracks (see Fig. 4.8). The very slow cooling rate formed suitable conditions for the development of erosional surfaces during one or several planation events.

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Fig. 4.10: possible t-T paths for individual samples based on Fission track single grain ages and the confined track length distribution as well as the (U-Th)/He ages of apatite and zircon; light grey paths: acceptable fit, dark grey: good fit, black line: best fitting path. With exception of APM 35-08 all models are based on (U-Th)/He data from zircon and apatite and Apatite Fission track data. Model 35-08 is only based on (U-Th)/He data.

Based on model data, conservative calculation of cooling rates to temperatures around 175°C (as illustrated in Figures 4.10 and 4.11) always yields rates below 5°C/Ma. For the temperature range of ca. 175°C (PRZZ) to ca. 65°C (PRZA), rates vary from around 2°C/Ma to 10°C/Ma. An exception of slow to moderate cooling rates is given by samples APM 34-08 and 49-08, which yield rates of 0.5 to 1.5 °C/Ma. These very low cooling rates, with the observation of older ages for cooling to temperatures of ca. 175°C (see above), strongly indicate a different thermal history of samples 34-08 and 49-08 in comparison to all other samples, at least for the cooling above the PAZA (Fig. 4.11). Very

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low cooling rates of <0.5°C/Ma (mostly even smaller) are calculated for all models in the temperature range of 65°C to around 30°C. Final cooling to the present day mean surface temperature of 17°C (Müller 1996) is difficult to define because model paths are less constrained in this temperature range. However, calculations based on modelling results demonstrate cooling rates in the range of 0.5 to 1.5°C/Ma for the cooling to surface temperatures.

In summary, all models show slow cooling of less than 0.5°C/Ma since at least 200 Ma (APM 37-08) and less than 0.15°C/Ma since at least 165 Ma (APM 34-08). Although cooling to surface temperatures is methodically only poorly constrained, it is not unlikely to have occurred between 80-40 Ma (Fig. 4.10).

Fig. 4.11: Cooling history for the San Luis profile based on individual t-T paths presented in Fig. 4.10.Horizontal error bars decipher the range of good fits ,the position of the marker represent position of best fit path (see Fig. 4.9). For a better illustration the individual samples are exhibited at slightly different temperatures. The approximate temperature ranges of the Partial Retention Zone for zircon (PRZZ), the Partial Annealing Zone for apatite (PAZA) as well as the Partial Retention Zone for apatite (PRZA) are shown as grey bars. The high age span of sample APM 35-08 is due to lack of Apatite Fission Track data (see Fig. 4.10).