Record of fluid flow history through fractured conglomerates, Lower Old Red Sandstone of central Scotland

Martin Baron and John Parnell
Scottish Journal of Geology, 40, 145-157, 1 November 2004, https://doi.org/10.1144/sjg40020145

Synopsis

Extensively fractured quartzite clasts in Lower Old Red Sandstone conglomerates adjacent to the Highland Boundary Fault Zone in central Scotland contain multiple populations of healed microfractures delimited by fluid inclusion planes. Microthermometric analysis indicates that the healing of microfractures oriented perpendicular to the trend of the Highland Boundary Fault Zone, which probably formed during Acadian overthrusting towards the SE along the fault zone in Middle Devonian times, involved moderate temperature (Th 102 to 238°C), low to moderate salinity (3 to 17 wt% NaCl eq.) aqueous fluids. The clasts also contain a second set of fluid inclusion planes oriented parallel to the Highland Boundary Fault Zone, which formed during extension associated with the intrusion of late Carboniferous quartz dolerite dykes. The healing of this set of microfractures involved low temperature (Th 58 to 155°C), low to moderate salinity (1 to 12 wt% NaCl eq.) aqueous fluids.

Introduction

Microfracture networks are developed in most fault zones. Open microfractures, which subsequently heal, leave planes of individual inclusions that have entrapped fluids present at the healing of the microfractures (fluid inclusion planes - FIPs). The orientation of open microfractures and FIPs in fault zones can provide valuable information concerning the palaeostress regime during brittle deformation (Wise 1964; Wang & Simmons 1978; Lespinasse & Pecher 1986; Kowallis et al. 1987). FIPs are extensional fractures (Mode I) that display no evidence of simple shear displacement. Under a triaxial stress regime mode I fractures form at orientations that are predominantly parallel to the maximum principal compressive stress axis (σ1) with the minimum principal stress axis (σ3) perpendicular to the fractures. Therefore the geometries of FIPs can provide useful information concerning the stress histories of weakly deformed rocks. Experimental work has shown that open microcracks in quartz will have geologically short lifetimes at elevated temperatures in the presence of suitable aqueous fluids (Smith & Evans 1984; Brantley et al. 1990). Therefore individual fluid inclusions within FIPs can be considered to provide information about the pressure, temperature and composition of the fluids present at the time of microfracturing (Pecher et al. 1985; Lespinasse & Cathelineau 1990; Lespinasse 1999). Since FIPs are reliable structural indicators they can also provide useful information about the geometry of fluid migration within a stratigraphical unit in relation to periods of deformation.

This study presents a combination of structural information on open microfractures and FIPs within Lower Old Red Sandstone (ORS) conglomerate quartzite clasts in the immediate vicinity of the Highland Boundary Fault Zone (HBFZ) in central Scotland along with petrographic and microthermometric data from individual fluid inclusions within FIPs.

Studies investigating the pressure, temperature and composition of fluids present during stages of microfracturing and regional deformation events have been restricted to granitoid rocks (e.g. Jang et al. 1989; Lespinasse & Cathelineau 1995; Boiron et al. 2001), quartz veins (e.g. De Alvarenga et al. 1990; Boullier & Roberts 1992; Meere 1995) and gneisses (e.g. O'Hara & Haak 1992; Boullier et al. 1997). Studies of FIPs in sedimentary rocks have been purely concerned with elucidating palaeostress regimes during low-level deformation (e.g. Laubach 1989; Onasch 1990; Eidelman & Reches 1992). The main aims of this study are: (1) to investigate whether FIPs within Lower ORS conglomerate clasts can provide any information concerning temporal changes of local stress domains in the rocks adjacent to the HBFZ and whether this provides any information about the late-stage evolution of the HBFZ; (2) to assess whether these FIPs can be used to provide information about properties of the fluids present during periods of deformation and regional fluid migration.

Regional geology

The HBFZ in central Scotland separates the Midland Valley, which consists predominantly of Palaeozoic vol- canics, clastics and limestones, from Dalradian meta- morphic rocks to the north (see Fig. 1). The HBFZ contains a series of discontinuous, fault-bound slivers of weakly metamorphosed rocks termed the Highland Border Complex (HBC). The HBC consists of a series of ultrabasic igneous rocks, hornblende-schists, basic volcanics, cherts, mudstones, sandstones and limestones containing fossils of Ordovician age (Henderson & Robertson 1982; Curry et al. 1984). The relationship between the Dalradian rocks and the HBC has been a matter of much debate in recent years (see Bluck et al. 1997). One view is that the Dalradian and HBC were deformed in a single orogenic event during early to middle Ordovician times (e.g. Johnson & Harris 1967; Tanner 1995), whereas others believe that the Dalradian was deformed during a late-Proterozoic event, which did not affect the HBC because it was well apart from the Dalradian terrane until they docked during Devonian times (e.g. Rogers et al. 1989; Bluck & Dempster 1991).

Fig. 1.
Fig. 1.

Simplified geological map of the region surrounding the HBFZ with sample localities.

The northern Midland Valley, immediately south of the HBFZ is predominantly composed of a series of Lower ORS lavas and sediments, which range in age from latest Silurian to early Devonian (Trewin & Thirlwall 2002). These rocks are folded into the asymmetrical Strathmore syncline and Sidlaw anticline, which have folded the strata to such an extent that the beds are often vertical to overturned adjacent to the HBFZ (Trewin & Thirlwall 2002). Upper ORS sandstones and conglomerates, which are late Devonian to earliest Carboniferous in age, overlie Lower ORS strata with an angular unconformity, with Middle ORS rocks being totally absent from the Midland Valley (Bluck 1978).

Fractured Lower ORS conglomerates

Du Toit (1905) first documented the existence of extensively fractured pebbles with a predominantly north-south fracture orientation in Lower ORS conglomerates at Loch Lomond (Fig. 2). Ramsay (1962, 1964) later documented similar oriented fractures in shattered pebbles from Lower ORS conglomerates adjacent to the HBFZ at various localities between Loch Lomond and North Esk.

Fig. 2.
Fig. 2.

Prominent NW-SE trending fractures within Lower ORS strata at Balmaha Pier, Loch Lomond, showing fracturing in quartzite clasts orientated approximately parallel to the length of the hammer head (width of view 1 m). Horizontal view.

The fractured pebbles are related to numerous shear zones which traverse the conglomerates varying from small clean cut faults with little offset to broader zones of brecciation with extensive slickensiding (Ramsay 1964). These shear zones all have orientations parallel or oblique to the trends of the conspicuous fractures within the Lower ORS conglomerates. The width of the deformation zone south of the HBFZ in which pebbles within Lower ORS conglomerates are fractured varies between 0.4 and 3.25 km (Ramsay 1964). Ramsay (1962) found that the direction of extensional fracturing at Loch Lomond and North Esk is NW-SE, normal to the HBFZ, while in the central region of Scotland the dominant orientation changes to NNW-SSE when the trend of the fault changes from NE-SW to ENE-WSW (Ramsay 1964). The direction, which corresponds to the maximum frequency of single cracks and shear zones, is therefore taken to be the direction in which the principal stress acted. This direction is NW-SE to NNW-SSE, which is almost normal to the orientation of the HBFZ (Ramsay 1962).

Methodology

Oriented samples of highly fractured Lower ORS conglomerate quartzite clasts were collected from seven localities (see Fig. 1). The orientations of FIPs were measured on the same sections used for microthermometric analysis using the methods of Boullier & Robert (1992). These sections were prepared as 150 |^m thick doubly polished wafers using standard techniques (Shepherd et al. 1985). Both the azimuth and dips of the FIPs were measured using a standard optical microscope with a calibrated micrometric focusing screw. By combining the projected widths of the individual FIPs with the precise thickness of the doubly polished wafers using the calibrated micrometric focusing screw it was possible to calculate dips of FIPs. The sections were cut at angles approximately perpendicular to the dominant orientation of the open microfractures, which means that any FIPs perpendicular or at high angles to the dominant orientation of the open microfractures would be missed.

The petrographies of individual fluid inclusion populations within FIPs were first examined at low magnifications using a standard optical microscope. Microthermometric analysis of individual fluid inclusions was carried out using a calibrated LINKAM TH-600 stage and measured with the aid of a video screen coupled to an optical microscope. Homogenization temperature (Th) measurements were determined using a heating rate of 10°C min−1. First ice melting temperature (Te) and final ice melting temperature (Tm) were determined using a heating rate of 1 ° C min −1.

The polished wafers were also examined using a cathodoluminescence detector attached to a JOEL 6400 electron microscope (SEM-CL) in order to establish more accurately the chronology and width of the FIPs. Operating beam current conditions were 30 to 70 fiA. with an acceleration voltage of between 15 and 20 kV. SEM-CL examination was performed after microthermometric analysis in order to prevent excessive overheating of the fluid inclusions in the samples.

SEM-CL of quartz displays different intensities from black to white that correspond to different levels of luminescence. The causes of luminesence in quartz are poorly understood. The main factors that are believed to control luminescence include trace element chemistry, crystallographic defects and temperature of precipitation. The most likely trace element substitution responsible for luminescence in quartz is aluminium (Watt et al. 1997).

Orientations and morphologies of FIPs within fractured Lower ORS conglomerate quartzite clasts

The fractured quartzite clasts contain abundant FIPs. Transmitted light and SEM-CL microscopy reveals that all the clasts display complex, superimposed networks of FIPs, indicating that they have experienced repeated microfracturing and healing. However, the orientations of FIPs from the different localities are surprisingly consistent. Figure 3 shows that the majority of the orientations of FIPs fall into two main groups. The first group (FIP set I) is parallel to sub-parallel to the dominant trend of the open microfractures with NW-SE to NNW-SSE orientations. The second group (FIP set II) is at high angles to FIP set I and typically trends approximately parallel to the HBFZ with NE-SW to ENE-WSW orientations. The majority of the FIPs are vertical to dipping at high angles (usually more than 80°). The fine-grained sandy matrix of the Lower ORS conglomerates has regularly been squeezed into the prominent NW-SE to NNW-SSE trending open fractures. This may explain why these open fractures were not shut during the later episode of deformation that formed FIP set II. The fact that the open, unhealed NW-SE to NNW-SSE fractures often exhibit evidence for slight shear displacement also means that the fractures may have been wedged open by irregularities on the fracture surface.

Fig. 3.
Fig. 3.

Contoured lower hemisphere equal-area plots of poles of FIPs hosted within fractured Lower ORS quartzite conglomerate clasts adjacent to the HBFZ. Contour interval is 1.0 sigma. Black arrows give the maximum principal stress σ1 for FIP set I, and white arrows give the maximum principal stress σ1 for FIP set II.

All of the samples contain approximately twice as many FIP set I as FIP set II. Both groups contain some FIPs that cross grain boundaries as well as other FIPs that terminate at grain boundaries (Fig. 4). FIP set II within quartzite clasts from Loch Lomond and Aberfoyle also terminate at FIP set I. Both sets of FIPs typically extend over the length of the samples (>20 mm).

Fig. 4.
Fig. 4.

Transmitted light images of (a) FIP set II cross-cutting prominent open NW-SE trending fracture (F) within a quartzite clast from Lower ORS conglomerates at Keltie Water (width of view 850 mm); (b) FIP set II cross-cutting FIP set I within a fractured quartzite clasts from Lower ORS conglomerates at N. Esk (width of view 425 mm).

Under SEM-CL the quartzite clasts have dull or no luminescence. Both sets of FIPs are typically brightly luminescent and have widths of between 5 and 10 μm (Fig. 5a). They are straight or slightly irregular in morphology and they often occur in clusters of up to ten FIPs, which are parallel to one another (Fig. 5b). SEM-CL is able to identify many FIPs that are not identified using transmitted light because only FIPs that contain a high number of fluid inclusions are distinguishable under transmitted light microscopy.

Fig. 5.
Fig. 5.

SEM-CL images of (a) multiple sets of FIPs within a fractured quartzite clast from the Lower ORS conglomerates at Aberfoyle (width of view 0.8 mm); (b) FIP set I occurring in approximately parallel sets of clusters within a fractured quartzite clast from the Lower ORS conglomerates at Hill of Alyth (width of view: 0.8 mm). SEM-CL reveals more precise information about the number, chronology and widths of FIPs than transmitted light microscopy.

Since the orientations of FIPs reflect the direction of the local maximum principal stress, FIP set I will have formed in a different stress regime to FIP set II. The relative timing of the formation of FIP set I and FIP set II is difficult to determine in most of the samples under transmitted light because they are extensional features and therefore the relative timing of the formation of FIPs cannot be determined from displacement relationships. However, some of the samples contain shear fractures orientated parallel to the open NW-SE to NNW-SSE trending microfractures and FIP set I. The orientation of these shear fractures suggests that they probably formed at the same time as the open NW-SE to NNW-SSE trending microfractures and FIP set I. When FIP set II crosses these shear fractures they are not displaced suggesting that FIP set II formed after the shear fractures and therefore also the open NW-SE to NNW-SSE trending microfractures and FIP set I. SEM-CL also shows that when FIP set I intersects FIP set II, the FIP set I is occasionally slightly laterally displaced (c. 10 to 50 |um), indicating that FIP set I formed before FIP set II.

It is unlikely that FIP set I was reoriented during the formation of FIP set II because the deformation was purely extensional. The lack of evidence for large-scale shear displacement along the HBFZ after Acadian deformation also makes it unlikely that either set of FIPs experienced significant reorientation since their formation.

Fluid inclusion data

Microthermometric analysis was only performed on fluid inclusions within FIPs sets I and II in order to avoid fluid inclusions that may have been inherited from an earlier stage in the history of the quartzite, such as within a metamorphic terrane or an earlier conglomerate.

Fluid inclusion petrography

Within the majority of the samples FIP set I contains more fluid inclusions suitable for microthermometric analysis than FIP set II. At room temperature individual fluid inclusions within FIPs are two-phase (liquid and vapour filled) with volumetric proportions of liquid ranging from approximately 90 to 98%. These inclusions typically possess flattened and sub-negative crystal shape morphologies and range in size from 2 to 12 μm.

Fluid inclusion microthermometry

A total of 959 measurements were recorded from 632 separate inclusions. A summary of the microthermometric data of all the phase changes observed is presented in Table 1 and Figures 6, 7, 8 and 9. Individual fluid inclusions within each set possess slightly different microthermometric properties. Te values of fluid inclusions within FIP set I range from −47.8 to −24.1 ° C. In contrast Te values of fluid inclusions within FIP set II range from −52.4 to −40.8°C.

View this table:
TABLE 1

Microthermometric data of fluid inclusions within FIP set I and FIP set II hosted within fractured Lower ORS quartzite conglomerate clasts adjacent to the HBFZ.

Fig. 6.
Fig. 6.

Final ice melting temperature versus first ice melting temperature of fluid inclusions within FIP set I (black circles) and FIP set II (white circles) hosted within fractured Lower ORS quartzite conglomerate clasts adjacent to the HBFZ.

Fig. 7.
Fig. 7.

Final ice melting temperature versus homogenization temperature of fluid inclusions within FIP set I (black circles) and FIP set II (white circles) hosted within fractured Lower ORS quartzite conglomerate clasts adjacent to the HBFZ at (a) Helensburgh SE, (b) Loch Lomond, (c) Aberfoyle, (d) Keltie Water, (e) Knock of Crieff, (f) Balwhyme, (g) N. Esk and (h) total data. Dashed lines enclose measurements from fluid inclusion within the same FIPs.

Fig. 8.
Fig. 8.

Drawing of fluid inclusions within FIP set I and FIP set II from Balwhyme. Fluid inclusions from different depths were projected into the plane of the drawing. Dashed lines represent traces of FIPs. Microthermometric measurements are presented next to the individual inclusions. Note that FIP set I and FIP set II display different microthermometric values indicating that they entrapped fluids with different temperatures and compositions.

Fig. 9.
Fig. 9.

Histograms of homogenization temperatures of fluid inclusions within FIPs from (a) Helensburgh SE, (b) Loch Lomond, (c) Aberfoyle, (d) Keltie Water, (e) Knock of Crieff, (f) Balwhyme and (g) N. Esk. White areas indicate measurements from FIP set I and black areas indicate measurements from FIP set II.

Fig. 10.
Fig. 10.

Isochore plots showing the possible ranges in temperature and pressure conditions during the formation of FIP set I and FIP set II calculated using the equations of state of Knight & Bodnar (1989) with the aid of the FLUIDS computer program (Bakker 2003). The highest and lowest density isochores from both sets of FIPs are displayed. Dashed lines represent possible ranges of hydrostatic fluid pressure conditions calculated using thermal gradients of 60 and 80 °C km -1 and fluid densities of 1 g cm-3. Shaded areas represent the estimated conditions of fluid entrapment.

The temperature at which ice first begins to melt during the heating of totally frozen fluid inclusions can be used to indicate what type of salts are present in solution by comparing the temperature to the eutectic temperature of pertinent aqueous-salt systems. For example, the eutectic temperature for H20-NaCl-CaCl2 is −55 °C, for H20-NaCl-MgCl2 it is −35 °C, for H20-NaCl-KCl it is −23.5°C and for H20-NaCl it is −21.2°C (Shepherd et al. 1985 and references therein). The different spread of Te measurements of fluid inclusions in FIP set I compared to FIP set II suggests that the fluids entrapped within the two sets of FIPs in each of the samples contained different types of salts. The relatively high Te values of fluid inclusions within FIP set I suggest that these fluids are dominated by H20 NaCl MgCl2 and H20 NaCl-KCl, whereas the low Tc values of fluid inclusions within FIP set II suggest that the fluids filling these inclusions are dominated by H20 NaCl CaCl2. However, it should be noted that this phase change in small inclusions is very subtle and the precise temperature at which it occurs is probably not well constrained.

Tm values of fluid inclusions within FIP set I range from −12.6 to −1.9°C, indicating that salinities are variable ranging from 3.23 to 16.53 wt% NaCl eq. (Bodnar 1993). Tm values of fluid inclusions within FIP set II range from —8.1 to −0.8°C, equating to salinities in the range of 1.40 to 11.81 wt% NaCl eq. (Bodnar 1993). Th values of fluid inclusions within FIP set I range from 102 to 238 °C, whereas Th values of fluid inclusions within FIP set II are slightly lower ranging from 58 to 155°C.

Discussion

Fluid characteristics during microfracturing

Fluid inclusions within some of the FIP set I contain variable salinities (see Figs. 7 and 8). The possibility that some of the higher salinity aqueous fluids could have been derived from magmatic or metamorphic processes seems unlikely because the fluids involved in mineralization in other areas of central Scotland that were derived from magmatic and metamorphic sources have H20-C02-salt compositions (e.g. Curtis et al. 1993; Craw & Chamberlain 1996). The saline composition precludes unaltered meteoric or marine waters as being the main source of these moderate salinity fluids. The moderate salinity aqueous composition is most typical of basinal brines (e.g. Walderhaug 1994) whose salinity is attributed to extensive water-rock interactions, dissolution of pre-existing evaporites or directly from bittern fluids. The lower salinity, aqueous compositions of some of the fluid inclusions is typical of surface-derived waters sourced from meteoric and/or marine origins.

Fluid salinities of fluid inclusions within FIP set II display more consistent values over the length of the fault zone.

Figures 7 and 8 shows that some FIPs contain variable salinities suggesting that fluid mixing may have occurred during microfracturing. The wide variation of Te values of fluid inclusions within both sets of FIPs also suggests that fluids contained variable Na, K, Mg and Ca contents. However, turbulent flow and mixing in a fracture a few micrometres in width is likely to be very limited, suggesting that another process may be responsible for the variable salinities in some of the FIPs. It is possible that the FIPs may have been partly reopened with some of the inclusions trapping fluid from the later event and some the earlier. In theory, a FIP would be a surface of mechanical weakness, which may be the position where a new fracture could form if the rock was in the same stress regime as when the original FIP formed. Another possibility is that the fluid inclusions interpreted as being in the same FIP are in fact from different FIPs. SEM-CL reveals that FIPs often occur in closely spaced clusters (see Fig. 5). Fluid inclusions that appear to be within the same FIP under transmitted light may in fact be within adjacent FIPs, especially if the fluid inclusions are at different depths within the sample.

At each locality Th values of fluid inclusions within both sets of FIPs display a considerable variation (see Figs. 8 and 9). There are a number of possible explanations for the spread of Th values. The most likely explanations for this spread are that there were pressure fluctuations during fluid entrapment, the inclusions trapped a mixture of high and low temperature fluids or the individual inclusions experienced different degrees of post-entrapment modification.

Large-scale, cyclic fluid pressure fluctuations are believed to occur in the vicinity of faults in response to tectonic activity (Sibson et al. 1988). Such fluctuations in fluid pressure could explain the considerable variation in Th values of fluid inclusions within FIPs. A fluid inclusion trapped at relatively low pressures would yield a lower Th value than a fluid inclusion trapped at relatively high pressures if the inclusions were entrapped at the same temperatures. It is therefore possible that fluid inclusions within FIPs that yield variable Th values were entrapped from a fluid of constant temperature, but variable pressure.

At all the localities the fluid inclusions within FIP set I yield higher average Th values than the fluid inclusions within FIP set II, which strongly suggests that the ambient temperatures were greater during the formation of FIP set I.

The large spread of Th values of fluid inclusions within both sets of FIPs may also be due to reequilibration of inclusions to lower densities after they had nucleated a vapour bubble. The inclusions within individual FIPs with higher Th values may have experienced higher degrees of re-equilibration. Laboratory experiments performed in order to determine how elevated temperatures and pressures re-equilibrate inclusions (Bodnar et al. 1989; Sterner & Bodnar 1989; Bakker & Jansen 1990) have shown that the degree of re-equilibration is dependent on the strength of the mineral host and to a lesser degree on the shape and size of the inclusion. Fluid inclusions re-equilibrate to lower densities by either inelastic volume change, leakage by dislocations or decrepitation due to internal overpressures. Bodnar et al. (1989) demonstrated that quartz- hosted inclusions less than 5 jim in diameter can withstand internal overpressures (i.e. the difference between the internal pressure of the fluid in the inclusions and the confining pressure acting upon the mineral host) of over 2 kbar. For lithostatic conditions, fluid pressures of 2 kbar correspond to burial depths of 8 to 9 km for these types of rocks. It is unlikely that these Lower ORS conglomerates were buried to such great depths because the prominent open fractures and FIP set I formed soon after the conglomerates were deposited during Middle Devonian times.

Tm and Th values can be used to construct isochores illustrating the variations in pressure-temperature conditions of fluid entrapment for individual inclusions. The intersection of these isochores and geothermobaric gradient provides a rough estimation of the pressure- temperature conditions during fluid entrapment. Pressure conditions were assumed to be hydrostatic because the hydrothermal fluids were migrating in well developed, interconnected open microfractures. Excessive fluid pressures are also thought to be unlikely because fracturing occurred soon after the conglomerates were deposited during a period of rapid uplift, a setting where regional pore fluid pressures are usually very low (e.g. Corbet & Bethke 1992). The precise thermal gradient is unknown, but from analogy with other ORS sequences in Scotland (e.g. Marshall 1998) it can be considered to be particularly high (c. 60-80 °C km - *) due to elevated upper crustal heat flow associated with Caledonian magmatism. Figure 10 shows that a pressure correction of between 5 and 27 ° C needs to be applied to Th values of fluid inclusions within FIP set I and a pressure correction of between 4 and 13 °C needs to be applied to Th values of fluids inclusions within FIP set II.

Age constraints and regional mineralization

Timing of FIP set I. We have argued that FIP set I is older than FIP set II. The temperature of the fluids entrapped within FIP set I was probably higher than that of fluids entrapped within FIP set II, suggesting that fracturing and rehealing of FIP set I and FIP set II most likely occurred during the progressive denudation of the fault zone.

The Lower ORS conglomerates have been tightly folded against the HBFZ. The overlying Upper ORS and younger strata above the Acadian unconformity are gently dipping and undeformed indicating that the most obvious deformation event these conglomerates have experienced was during the Acadian orogeny. The prominent open NW-SE to NNW-SSE trending microfractures do not extend into the Upper ORS strata showing convincing evidence that they formed during the Acadian orogeny in Middle Devonian times. Under high magnifications the open NW-SE to NNW-SSE trending microfractures commonly occur surrounded by numerous parallel trending FIPs indicating that there is a gradation between the open microfracture set and FIP set I. The open NW-SE to NNW-SSE trending microfractures have probably not healed because they were wider than the fractures that did heal or they never contained solutions suitable to heal them. The similar orientation of the FIP set I and the prominent NW-SE to NNW-SSE open microfractures means that they were produced in the same stress regime and therefore are most likely to be broadly coeval. Experimental data of Smith & Evans (1984) showed that at a temperature of 200 °C, which is similar to the temperature at which FIP set I healed, microfractures heal in a matter of hours. This means that the fluid inclusions entrapped within the FIPs contain fluids that were present at the time of fracturing.

These lines of evidence indicate that FIP set I formed during overthrusting towards the SE along the HBFZ during Mid-Devonian times when the maximum principal compressive stress was at NW-SE to NNW-SSE orientations, perpendicular to the trace of the HBFZ (Ramsay 1964). This widespread Caledonian structural episode is also recognized south of the Iapetus suture in the Appalachians of North America and South Wales where it is also delimited by a major unconformity (Bluck 1984).

Timing of FIP set II. The contrasting orientation of FIP set II strongly suggests that it formed under a different stress regime from FIP set I. Additionally microthermometric results from fluid inclusions within

FIP set I and FIP set II indicate that they trapped fluids with contrasting compositions. This indicates that FIP set II formed during a period of tectonism associated with the HBFZ post-Mid-Devonian times. The ENE-WSW to NE-SW orientations of FIP set II indicate that fracturing occurred when the minimum principal stress was at WSW-ENE to NW-SE, perpendicular to the trace of the HBFZ. Given the orientation of FIP set II and the regional tectonic history of the Northern Midland Valley it is most likely that FIP set II formed during Late Carboniferous to early Permian times.

The orientations of FIP set II are also similar to trends of late Carboniferous quartz dolerite dykes in the region surrounding the HBFZ. Although the late Carboniferous quartz dolerite dykes exhibit a dominant east-west orientation, the HBFZ causes the dykes to deflect into trends approximately parallel to the fault (see Fig. 11). The fact that the trends of some of the dykes follow the trace of the fault indicates that the HBFZ acted as a conduit for magmatism during the late Carboniferous. The intrusion of this dyke suite occurred as a result of lithospheric fracturing across a zone 300 km wide as a consequence of the oblique collision of the Variscan-Alleghenian orogeny, in late Carboniferous-early Permian times following the closure of the mid-European Rheic ocean (Lorenz 1976; Smythe et al. 1995). This phase of extensional stress also resulted in major rifting in the Oslo Graben as well as other basins in the North Sea (Russell & Smythe 1983).

Fig. 11.
Fig. 11.

Orientations of late Carboniferous quartz dolerite dykes in central Scotland showing the deviation of the dykes from an arcuate east-west regional trend to trends that trail approximately parallel to the HBFZ in areas adjacent to the fault (adapted from Smythe et al. 1995).

There are also numerous lines of evidence for a regionally extensive episode of mineralization occurring in central Scotland during late Carboniferous times. Palaeomagnetic studies have shown that iron oxide minerals within highly altered serpentinites of the HBC were precipitated during late Carboniferous to early Permian times (Elmore et al. 2002). Fluid inclusion microthermometric studies of inclusions within carbonate and quartz minerals associated with this late Carboniferous to early Permian episode of mineralization involved fluids of similar temperature and salinity (Elmore et al. 2002). Parnell et al. (2000a) also recognized high levels of heat flow in the vicinity of the HBFZ during late Carboniferous to early Permian using clay poly type studies. Kaolinite minerals associated with mineralization along the whole length of the HBFZ in central Scotland were transformed to the high temperature kaolin polytype, dickite, in areas comparable with the spatial distribution of the late Carboniferous quartz dolerite dykes (Parnell et al. 2000a). There are also numerous occurrences of hydrothermal Ca-Ba-Pb-Zn mineralization in association with late Carboniferous quartz dolerite dykes in central Scotland (Ineson & Mitchell 1974; Gallagher 1964). These lines of evidence attest to the presence of a regionally extensive mineralization event in the area in and around the HBFZ of central Scotland. It is noteworthy that a stable isotope study of fluid infiltration associated with the HBFZ, near Stonehaven (Masters et al. 2000) recognized the presence of a late-stage post-Caledonian episode of mineralization, at temperatures of approximately 150°C. Elmore et al. (2002) recorded vein-filling carbonate minerals with similar isotopic signatures at other localities along the HBFZ, suggesting that this late-stage post-Caledonian episode of mineralization was widespread. There is also evidence from the location of base metal sulphides in association with the continuation of the HBFZ in the north of Ireland that the HBFZ acted as a major conduit for saline mineralizing solutions during Carboniferous times (Wilkinson et al. 1999; Hitzman 1986).

There are also numerous occurrences of reddened dolomitic breccia veins in the Dalradian rocks of SW Scotland that formed during late Carboniferous times (Parnell et al. 2000b). Fluid inclusion studies from these dolomitic breccia veins yielded very similar temperatures and salinities to the fluid inclusions within FIP set II.

The salinities of the fluids entrapped within both sets of FIPs are only low to moderate. These relatively low values are typical of connate waters whose salinities can be attributed to solely water-rock interactions.

This study has shown how the geometry of the fracture permeability and hence fluid circulation patterns within the Lower ORS conglomerates in the HBFZ changed with time. Fracture permeability in the fault zone during Mid-Devonian times was dominated by NW-SE to NNW-SSE trends, perpendicular to the HBFZ, due to the fact that the HBFZ was subjected to compression during the Acadian orogeny. In contrast, during Carboniferous times, when the regional stress regime was different from that of Mid-Devonian times, the dominant fracture permeability was at NE-SW to ENE-WSW orientations, approximately parallel to the HBFZ. This later phase of fracturing formed in response to regional extension associated with the intrusion of a late Carboniferous quartz dolerite dyke suite.

Conclusions

Combined structural, petrographic and fluid inclusion microthermometric analysis of FIPs within fractured Lower ORS conglomerate quartzite clasts have recorded temporal changes of local stress domains in the HBFZ as well as the physical and chemical characteristics of the ambient fluids present during separate episodes of tectonism. The main findings of this study are as follows:

  1. Fractured Lower ORS quartzite conglomerate clasts contain two dominant sets of FIPs (FIP set I and FIP set II) that display consistent orientations at localities over 100 km apart.

  2. FIP set I possesses azimuths dominantly of NW-SE to NNW-SSE orientations that are parallel and sub-parallel to the major trends of open microfractures. The formation of FIP set I occurred when the maximum principal compressive stress was at NW-SE to NNW-SSE orientations, perpendicular to the trace of the HBFZ. FIP set II possesses azimuths dominantly of NE-SW to ENE-WSW orientations that formed when the maximum principal compressive stress was at NE-SW to ENE-WSW orientations and the minimum principal stress was at NW-SE to NNW-SSE orientations, perpendicular to the trace of the HBFZ.

  3. Microthermometric analysis of individual fluid inclusions within FIPs indicate that the ambient fluids involved in healing FIP set I and FIP set II possessed contrasting compositions and temperatures.

  4. The formation of FIP set I that occurred during Acadian overthrusting towards the SE along the HBFZ during Middle Devonian times, involved moderate temperature (Th 102 to 238 °C) low to moderate salinity (3 to 17wt% NaCl eq.) aqueous fluids.

  5. The formation of FIP set II probably occurred during regional extension associated with the intrusion of late Carboniferous quartz dolerite dyke suite. It involved low temperature (Th 58 to 155°C) low to moderate salinity (1 to 12wt% NaCl eq.) aqueous fluids.

  6. The geometry of palaeofracture permeability and hence fluid circulation patterns within the Lower ORS conglomerates in the HBFZ changed with time due to variation in the regional stress regime. Fracture permeability during Mid-Devonian times was dominated by NW-SE to NNW-SSE trends, perpendicular to the HBFZ, whereas during late Carboniferous to early Permian times the dominant fracture permeability trended NE-SW to ENE- WSW, approximately parallel to the HBFZ.

Acknowledgements

We are very grateful for the critical reviews of the manuscript by P. Murphy and an anonymous reviewer, and editing of R. Ellam. Work for this study was done whilst the senior author (M.B.) was in receipt of a PhD grant provided by NIDevR, ESF and the Queen's University of Belfast.

References

  1. ↵
    1. BAKKER, R. J.
    2003. Package FLUIDS 1. Computer programs for analysis of fluid inclusion data and for modelling bulk fluid properties. Chemical Geology, 194, 3–23.
    OpenUrlCrossRefWeb of Science
  2. ↵
    1. BAKKER, R. J. &
    2. JANSEN, J. B. H.
    1990. Preferential water leakage from fluid inclusions by means of mobile dislocations. Nature, 345, 58–60.
    OpenUrlCrossRefWeb of Science
  3. ↵
    1. BLUCK, B. J.
    1978. Sedimentation in a late orogenic basin: the Old Red Sandstone, Scotland. In Balance, P. F. & Reading, H. G. (eds.) Crustal evolution in northwestern Britain and adjacent regions. Geological Journal Special Issue, 10, 249–278.
    OpenUrl
  4. ↵
    1. BLUCK, B. J.
    1984. Pre-Carboniferous history of the Midland Valley of Scotland. Transactions of the Royal Society of Edinburgh: Earth Sciences, 75, 275–295.
    OpenUrlCrossRef
  5. ↵
    1. BLUCK, B. J. &
    2. DEMPSTER, T. J.
    1991. Exotic metamorphic terranes in the Caledonides: Tectonic history of the Dalradian block. Geology, 19, 1133–1136.
    OpenUrlAbstract/FREE Full Text
  6. ↵
    1. BLUCK, B. J.,
    2. INGHAM, J. K. &
    3. TANNER, P. W. G.
    1997. The Highland Border controversy: a discussion and reply of “New evidence that the Lower Cambrian Leny Limestone at Callander, Perthshire, belongs to the Dalradian Supergroup, and a reassessment of the ‘exotic’ status of the Highland Border Complex”. Geological Magazine, 134, 563–570.
    OpenUrlAbstract
  7. ↵
    1. BODNAR, R. J.
    1993. Revised equation and table for determining the freezing point depression of H2O-NaCl solutions. Geochimica Cosmochimica Acta, 57, 683–684.
    OpenUrlCrossRef
  8. ↵
    1. BODNAR, R. J.,
    2. BINNS, P. R. &
    3. HALL, D. L.
    1989. Synthetic fluid inclusions–VI. Quantitative evaluation of the decrepitation behaviour of fluid inclusions in quartz at one atmosphere confining pressure. Journal of Metamorphic Geology, 7. 229–242.
    OpenUrlCrossRefWeb of Science
  9. ↵
    1. BOIRON, M. C.,
    2. BARAKAT, A.,
    3. CATHELINEAU, M.,
    4. BANKS, D. A.,
    5. DURISOVA, J. &
    6. MAORAVEK, P.
    2001. Geometry and P-V-T-X conditions of microfissural ore fluid migration: the Mokrsko gold deposit (Bohemia). Chemical Geology, 173, 207–225.
    OpenUrlCrossRefWeb of Science
  10. ↵
    1. BOULLIER, A. M. &
    2. ROBERT, F.
    1992. Palaeoseismic events recorded in the Archaean gold-quartz vein networks, Val d’Or, Albitibi, Quebec, Canada. Journal of Structural Geology, 14, 161–179.
    OpenUrlCrossRefWeb of Science
  11. ↵
    1. BOULLIER, A. M.,
    2. FIRDAOUS, K. &
    3. BOUDIER, F.
    1997. Fluid circulation related to deformation in the Zabargad gneisses (Red Sea rift). Tectonophvsics, 279, 281–302.
    OpenUrlCrossRef
  12. ↵
    1. BRANTLEY, S. L.,
    2. EVANS, B.,
    3. HICKMAN, S. H. &
    4. CRERAR, D. A.
    1990. Healing of microcracks in quartz: implications for fluid flow. Geology, 18, 136–139.
    OpenUrlAbstract/FREE Full Text
  13. ↵
    1. CORBET, T. F. &
    2. BETHKE, C. M.
    1992. Disequilibrium fluid pressures and groundwater flow in the Western Canada Sedimentary Basin. Journal of Geophysical Research, 97, 7203–7217.
    OpenUrlCrossRef
  14. ↵
    1. CRAW, D. &
    2. CHAMBERLAIN, C. P.
    1996. Meteoric incursion and oxygen fronts in the Dalradian metamorphic belt, southwest Scotland: a new hypothesis for regional gold mobility. Minerahum Deposita, 31, 365–373.
    OpenUrlCrossRef
  15. ↵
    1. CURRY, G. B.,
    2. BLUCK, B. J.,
    3. BURTON, C. J.,
    4. INGHAM, J. K.,
    5. SIVETER, D. J. &
    6. WILLIAMS, A.
    1984. Age. evolution and tectonic history of the Highland Border Complex, Scotland. Transactions of the Royal Society of Edinburgh: Earth Science, 75, 113–133.
    OpenUrlCrossRef
  16. ↵
    1. CURTIS, S. F.,
    2. PATTRICK, R. A. D.,
    3. JENKIN, G. R. T.,
    4. FALLICK, A. E.,
    5. BOYCE, A. J. &
    6. TREAGUS, J. E.
    1993. Fluid inclusion and stable isotope study of fault-related mineralization in the Tyndrum area, Scotland. Transactions of the Institution of Mining and Metallurgy (Section B Applied Earth Sciences), 102, B39–B47.
    OpenUrl
  17. ↵
    1. DE ALVARENGA, C. J. S.,
    2. CATHELINEAU, N. &
    3. DUBESSY, J.
    1990. Chronology and orientation of N2-CH4, H2 O-CO2 and H2O-rich fluid-inclusion trails in intrametamorphic quartz veins from the Cuiaba gold deposit, Brazil. Mineralogical Magazine, 54, 245–255.
    OpenUrlCrossRef
  18. ↵
    1. DU TOIT, A. L.
    1905. The Lower Old Red Sandstone rocks of the Balmaha-Aberfoyle region. Transactions of the Geological Society of Edinburgh, 8, 315–325.
    OpenUrlCrossRef
  19. ↵
    1. EIDELMAN, A. &
    2. RECHES, Z.
    1992. Fractured pebbles-A new stress indicator. Geology, 20, 307–310.
    OpenUrlAbstract/FREE Full Text
  20. ↵
    1. ELMORE, R. D.,
    2. PARNELL, J.,
    3. ENGEL, M.,
    4. BARON, M.,
    5. WOODS, S.,
    6. ABRAHAM, M. &
    7. DAVIDSON, M.
    2002. Palaeomagnetic dating of fluid-flow events in dolomitized rocks along the Highland Boundary Fault, central Scotland. Geofluids, 2, 299–314.
    OpenUrlCrossRefWeb of Science
  21. ↵
    1. GALLAGHER, M. J.
    1964. Rock alteration in some mineralized basic dykes in Britain. Transactions of the Institute of Mining and Metallurgy 73B, 825–840.
    OpenUrl
  22. ↵
    1. HENDERSON, W. G. &
    2. ROBERTSON, A. H. F.
    1982. The Highland Border rocks and their relation to marginal basin development in the Scottish Caledonides. Journal of the Geological Society, London, 139, 433–450.
    OpenUrlAbstract/FREE Full Text
  23. ↵
    1. HITZMAN, M. W.
    1986. The geology of the Abbeytown mine, Co. Sligo, Ireland. In Andrew, C. J., Crowe, R. W. A., Finlay, S., Pennell, W. H. & Pyne, J. F. (eds.) Geology and genesis of mineral deposits in Ireland, Irish Association of Economic Geology, 341–354.
  24. ↵
    1. INESON, P. R. &
    2. MITCHELL, J. G.
    1974. K-Arisotopic age determinations from some Scottish mineral localities. Transactions of the Institute of Mining and Metallurgy 83B, 13–18.
    OpenUrl
  25. ↵
    1. JANG, J.,
    2. WANG, H. F.,
    3. REN, X. &
    4. KOWALLIS, B. J.
    1989. Precambrian paleostress from microcracks and fluid inclusions in the Wolf River batholith of central Wisconsin. Geological Society of America Bulletin, 101, 1457–1464.
    OpenUrlAbstract/FREE Full Text
  26. ↵
    1. JOHNSON, M. R. W. &
    2. HARRIS, A. L.
    1967. Dalradian-?Arenig relations in parts of the Highland Border, Scotland, and their significance in the chronology of the Caledonian orogeny. Scottish Journal of Geology, 3, 1–16.
    OpenUrlAbstract/FREE Full Text
  27. ↵
    1. KNIGHT, C. L. &
    2. BODNAR, R. J.
    1989. Synthetic fluid inclusions: IX. Critical PVTX properties of H2O-NaCl solutions. Geochimica Cosmochimica Acta, 53, 3–8.
    OpenUrlCrossRefWeb of Science
  28. ↵
    1. KOWALLIS, B. J.,
    2. WANG, H. F. &
    3. JANG, B.
    1987. Healed microcrack orientations in granite from Illonois borehole UPH-3 and their relationship to the rock’s stress history. Tectonophysics, 135, 297–306.
    OpenUrlCrossRefWeb of Science
  29. ↵
    1. LAUBACH, S. E.
    1989. Palaeostress directions from the preferred orientation of closed microfractures (fluid-inclusion planes) in sandstones, East Texas basin, U.S.A.. Jowned of Structural Geology, 11, 603–611.
    OpenUrlCrossRef
  30. ↵
    1. LESPINASSE, M.
    1999. Are fluid inclusion planes useful in structural geology?. Journal of Structural Geology, 21, 1237–1243.
    OpenUrlCrossRefWeb of Science
  31. ↵
    1. LESPINASSE, M. &
    2. CATHELINEAU, M.
    1990. Fluid percolation in a fault zone: a study of fluid inclusion planes in the St Sylvestre granite, northwest Massif Central, France. Tectonophysics, 184, 173–187.
    OpenUrlCrossRefWeb of Science
  32. ↵
    1. LESPINASSE, M. &
    2. CATHELINEAU, M.
    1995. Palaeostress magnitude determination using fault slip and fluid inclusion planes data. Journal of Geophysical Research, 100, 3895–3904.
    OpenUrlCrossRef
  33. ↵
    1. LESPINASSE, M. &
    2. PECHER, A.
    1986. Microfracturing and regional stress field: a study of the preferred orientations of fluid-inclusion planes in a granite from the Massif Central, France. Journal of Structural Geology, 8, 169–180.
    OpenUrlCrossRefWeb of Science
  34. ↵
    1. LORENZ, V.
    1976. Formation of the Hercynian subplates, possible causes and consequences. Nature, 262, 374–377.
    OpenUrlCrossRef
  35. ↵
    1. MARSHALL, J. E. A.
    1998. The recognition of multiple hydrocarbon generation episodes: an example from Devonian lacustrine sedimentary rocks in the Inner Moray Firth, Scotland. Journal of the Geological Society, London, 155, 335–352.
    OpenUrlAbstract/FREE Full Text
  36. ↵
    1. MASTERS, R. L.,
    2. ARGUE, J. J. &
    3. RYE, D. M.
    2000. An oxygen and carbon isotopic study of multiple episodes of fluid flow in the Dalradian and Highland Border Complex, Stonehaven, Scotland. Journal of the Geological Society of London, 157, 367–379.
    OpenUrlAbstract/FREE Full Text
  37. ↵
    1. MEERE, P. A.
    1995. High and Low density fluids in a quartz vein from the Irish Variscides. Journal of Structural Geology, 17, 435–446.
    OpenUrlCrossRefWeb of Science
  38. ↵
    1. ONASCH, C. M.
    1990. Microfractures and their role in deformation of a quartz arenite from the central Appalachian foreland. Journal of Structural Geology, 12, 883–894.
    OpenUrlCrossRefWeb of Science
    1. O’HARA, K. &
    2. HAAK, A.
    1992. A fluid inclusion study of fluid pressure and salinity variations in the footwall of the Rector Branch thrust, North Carolina, U.S.A.. Journal of Structural Geology, 14, 579–589.
    OpenUrlCrossRefWeb of Science
  40. ↵
    1. PARNELL, J.,
    2. BARON, M. &
    3. BOYCE, A.
    2000a. Controls on kaolinite and dickite distribution, Highland Boundary Fault zone, Scotland and Northern Ireland. Journal of the Geological Society of London, 157, 635–640.
    OpenUrlAbstract/FREE Full Text
  41. ↵
    1. PARNELL, J.,
    2. BARON, M.,
    3. DAVIDSON, M.,
    4. ELMORE, D. &
    5. ENGEL, M.
    2000b. Dolomitic breccia veins as evidence for extension and fluid flow in the Dalradian of Argyll. Geological Magazine, 137, 447–462.
    OpenUrlAbstract/FREE Full Text
  42. ↵
    1. PECHER, A.,
    2. LESPINASSE, M. &
    3. LEROY, J.
    1985. Relations between fluid inclusion trails and regional stress field: a tool for fluid chronology-An example of an intragranitic uranium ore deposit (northwest Massif Central, France). Lithos, 18, 229–237.
    OpenUrlCrossRefWeb of Science
  43. ↵
    1. RAMSAY, D.
    1962. The Highland Boundary Fault: reverse or wrench fault? Nature, 195, 1190–1191.
    OpenUrlWeb of Science
  44. ↵
    1. RAMSAY, D. M.
    1964. Deformation of pebbles in Lower Old Red Sandstone conglomerates adjacent to the Highland Boundary Fault? Geological Magazine, 101, 228–248.
    OpenUrlAbstract
  45. ↵
    1. ROGERS, G.,
    2. DEMPSTER, T. J.,
    3. BLUCK, B. J. &
    4. TANNER, P. W. G.
    1989. A high precision U-Pb age for the Ben Vuirich granite: implications for the evolution of the Scottish Dalradian Supergroup. Journal of the Geological Society, London, 146, 789–798.
    OpenUrlAbstract/FREE Full Text
  46. ↵
    1. RUSSELL, M. J. &
    2. SMYTHE, D. K.
    1983. Origin of the Oslo Graben in relation to the Hercynian-Alleghenian orogeny and lithospheric rifting in the North Atlantic. Tectonophysics, 94, 457–472.
    OpenUrlCrossRefWeb of Science
  47. ↵
    1. SHEPHERD, T. J.,
    2. RANKIN, A. H. &
    3. ALDERTON, D. H. M.
    1985. A Practical Guide to Fluid Inclusion Studies, Blackie, London.
  48. ↵
    1. SIBSON, R. H.,
    2. ROBERT, F. &
    3. POULSEN, K. H.
    1988. High-angle reverse faults, fluid-pressure cycling, and mesothermal gold-quartz deposits. Geology, 16, 551–555.
    OpenUrlAbstract/FREE Full Text
  49. ↵
    1. SMITH, D. L. &
    2. EVANS, B.
    1984. Diffusional crack healing in quartz. Journal of Geophysical Research, 89, 4125–4135.
    OpenUrlCrossRef
  50. ↵
    1. SMYTHE, D. K.,
    2. RUSSELL, M. J. &
    3. SKUCE, A. G.
    1995. Intra-continental rifting inferred from the major late Carboniferous quartz-dolerite dyke swarm of NW Europe. Scottish Journal of Geology, 31, 151–162.
    OpenUrlAbstract/FREE Full Text
  51. ↵
    1. STERNER, S. M. &
    2. BODNAR, R. J.
    1989. Synthetic fluid inclusions-VII. Re-equilibration of fluid inclusions in quartz during laboratory-simulated metamorphic burial and uplift. Journal of Metamorphic Geology, 7, 243–260.
    OpenUrlCrossRefWeb of Science
  52. ↵
    1. TANNER, P. W. G.
    1995. New evidence that the Lower Cambrian Leny Limestone at Callander, Perthshire, belongs to the Dalradian Supergroup, and a reassessment of the ‘exotic’ status of the Highland Border Complex. Geological Magazine, 132, 473–483.
    OpenUrlAbstract
  53. ↵
    1. TREWIN, N. H. &
    2. THIRLWALL, M. F.
    2002. Old Red Sandstone. In Trewin, N. H. (ed.) The Geology of Scotland, 4th edn, The Geological Society, London, 213–249.
  54. ↵
    1. WALDERHAUG, O.
    1994. Temperatures of quartz cementation in Jurassic sandstones from the Norwegian continental shelf-evidence from fluid inclusions. Journal of Sedimentary Research, A64, 311–323.
    OpenUrlAbstract/FREE Full Text
  55. ↵
    1. WANG, H. F. &
    2. SIMMONS, G.
    1978. Microcracks in crystalline rocks from 5.3 km depth in the Michigan basin. Journal of Geophysical Research, 83, 5849–5856.
    OpenUrlCrossRef
  56. ↵
    1. WATT, G. R.,
    2. WRIGHT, P.,
    3. GALLOWAY, S. &
    4. MCLEAN, C.
    1997. Cathodoluminescence and trace element zoning in quartz phenocrysts and xeoncrysts. Geochimica et Cosmochimica Acta, 61, 4337–4348.
    OpenUrlCrossRefWeb of Science
  57. ↵
    1. WILKINSON, J. J.,
    2. BOYCE, A. J.,
    3. EARLS, G. &
    4. FALLICK, A. E.
    1999. Gold remobilization by low-temperature brines: Evidence from the Curraghinalt gold deposits, Northern Ireland. Economic Geology, 94, 289–296.
    OpenUrlAbstract/FREE Full Text
  58. ↵
    1. WISE, D. U.
    1964. Microjointing in basement, Middle Rocky Mountains of Montana and Wyoming. Bulletin of the Geological Society of America, 75, 287–306.
    OpenUrlAbstract/FREE Full Text
Back to top
Alerts
Citation tools
Permissions
View PDF
Email to
Print
Download PPT
Bookmark this article