Local-scale folds within the Mississippian Madison Group of the frontal Montana Rockies preserve pre- and synfolding remagnetization data. Paleomagnetic results display inclinations of ∼70°, in contrast to the expected shallower directions for North American Mississippian rocks. The magnetization is chemical in origin, preserved in superparamagnetic to single-domain magnetite grains from fluid activity. Magnetic intensity results in the study area suggest that mineralization was more prevalent in the interior of the fold-and-thrust belt and diminished toward the east, resulting in lower intensities from less magnetite growth in the very frontal portions of the belt into the foreland. Fold test results of individual folds show syn- and prefolding remagnetizations as a function of location across the belt, with synfolding results in more westerly locations and prefolding results in the most frontal folds of the belt. By comparing our synfolding results with previously determined deformation ages for the Rocky Mountains, an Eocene (53.6 Ma) age for the remagnetization can be assigned. Based on the relative timing of remagnetization, a spatial pattern of folding in the study area is revealed. Major folding commenced (i.e., synfolding magnetization) during an Eocene remagnetization event, while the most frontal portion remained undeformed (i.e., prefolding magnetization) and was subsequently folded after regional remagnetization.


Remagnetization of carbonates is globally extensive (McCabe and Elmore, 1989; Jackson and Swanson-Hysell, 2012; Van der Voo and Torsvik, 2012). It is generally inferred that carbonate remagnetizations are of chemical origin, given that alternative mechanisms (thermal or viscous) can be rejected based on predicted burial temperatures (Elmore et al., 2012; McCabe and Elmore, 1989; Font et al., 2012; Zegers et al., 2003). Chemical carbonate remagnetization can be the result of the growth of magnetite to a stable single-domain state, preserving a paleomagnetic remanence. Several mechanisms have been proposed as the catalyst for this pervasive chemical remagnetization in carbonates, including organic matter maturation, illitization, orogenic fluid movement, and hydrocarbon migration (Katz et al., 1998). Of these mechanisms, two ideas for the facilitation of chemical remagnetization seem to dominate the literature, orogenic fluid movement and illitization. In these cases, magnetite authigenesis occurs during the interaction with an orogenic brine, when iron (Fe) is released during the process of converting smectite to illite, or as a combination of both (Elmore et al., 2012; Evans et al., 2000; Lewchuk et al., 2003).

Evans et al. (2012) used isotope data as support for a connection between chemically remagnetized rocks and orogenic fluid movement. Where the rocks were remagnetized, the isotope data revealed 87Sr/86Sr values that were more radiogenic than coeval seawater, suggesting that the area had experienced diagenetic fluid movement. Where the rocks were not remagnetized, the 87Sr/86Sr fell within the range of coeval seawater values, which would suggest the rocks were unaltered by outside fluids.

Alternatively, during illitization, smectite will transform to illite as temperatures increase with burial (Altaner and Ylagan, 1997). During illitization, iron is released, and this iron can enable magnetite growth (Hirt et al., 1993). Woods et al. (2002) and Gill et al. (2002) supported illitization-induced remagnetization through a presence-absence test. Both studies looked at sedimentary rocks in which the occurrence of illite, from clay diagenesis (illitization), was associated with chemically remagnetized rocks, whereas the presence of smectite was associated with primary and/or weaker secondary magnetizations (Gill et al., 2002).

The timing of remagnetization in Mississippian carbonates from NW Montana is examined in this study in order to constrain the paleomagnetic history in the study area and its relation with deformation. Prior studies in the region have mostly focused on regional sampling to examine the mechanism and relative timing of magnetization (Eldredge and Van der Voo, 1988; Gill et al., 2002; Elliott et al., 2006; O’Brien et al., 2006, 2007). This study targeted individual folds to test for synfolding remagnetization and spatial variation of magnetic intensities in the deformation belt in terms of the relative timing of acquisition.


The North American Cordillera extends from Canada to southern Mexico and was formed by Sevier and Laramide orogen-style deformation from the mid-Mesozoic to Eocene (Burchfiel et al., 1992; Dickinson, 2004). The portion of the Sevier orogen in NW Montana crops out from Glacier-Waterton National Park on the Canadian border to the Helena Salient in west-central Montana. This study focused on the very frontal segment of the Rocky Mountains in NW Montana. The frontal Rockies in the study region trend roughly N-S and were previously delineated into four subbelts with thrusts and associated folds in each section (Mudge, 1970, 1982). K-Ar dates from bentonites and structural evidence in the Montana Rockies indicate that deformation occurred ∼72–54 m.y. ago (Elliott et al., 2006; Hoffman et al., 1976; McMannis, 1965).

Samples were collected from the frontal portion of the Montana Rockies, where N-S–trending, eastward-verging folds and associated imbricate thrust faults sole into a regional décollement (Mudge, 1982). The regional décollement in the area dips to the west and progresses up section from Mesoproterozoic rocks in the west to Cretaceous rocks in the east (Sears, 2001). In the north, deformation began with emplacement of the Mesoproterozoic Belt-Purcell Supergroup over Cretaceous sedimentary rocks by the Lewis thrust fault (Willis, 1902; Sears et al., 2005). Within the frontal segment of the belt, deformation resulted in ∼20 large imbricate thrusts that merge into a décollement in Cambrian shale (Mudge, 1982), including the Allan, Palmer, Beaver, Norwegian, French, Home, and Diversion thrusts (Lageson, 1987). These thrusts and folds verge to the east and place Mississippian carbonates over younger Cretaceous sandstones and shales (Mudge, 1970). Later modification of the Rockies in this area included Basin and Range extension with reactivation of thrust ramps as normal faults (Fuentes et al., 2011).

Mesoproterozoic rocks dominate the Montana Rockies except on the very eastern edge, where Paleozoic and Mesozoic rocks crop out (McMannis, 1965). The paleomagnetic target unit of this study was the Madison Group, consisting of limestone and dolomite that are typically thrusted on Cretaceous shales throughout the eastern (frontal) edge of the frontal Rockies. The Madison Group was deposited on a shallow platform during the Mississippian (359–323 Ma; Smith et al., 2004). It is estimated that the Madison unit is ∼500 m thick in the frontal Rockies and thins to the east, toward the foreland (Deiss, 1943; Mudge, 1970).

The Madison Group is an unconformity-bounded carbonate unit with two internal composite sequences (Smith et al., 2004). It is subdivided into the Allan Mountain and Castle Reef Limestones in the frontal Rockies, which are equivalent to the Lodgepole and Mission Canyon Limestones in the foreland (Mudge et al., 1962). The Madison Group has been thrusted on top of the Cretaceous Kootenai Formation, and both the Sun River and Teton Canyon transversely cut across the range, allowing for sampling of the folded Madison Group (Lageson, 1987).



The Mississippian Madison Group was the target of this study because access to local-scale folds is possible (Fig. 1). Two sites were also collected from the Cretaceous Colorado Group (HS1 and HS2), but these resulted in spurious demagnetization patterns and will not be discussed further. From the Madison Group, 17 sites were sampled: 13 from five local-scale folds (including three fold hinges) and four individual sites. One of the individual sites, SR1, was collected from a small exposure of the Madison Group in the foreland, also called the Sweetgrass Arch. The Sweetgrass Arch lies to the east of the Rockies in a relatively flat-lying, broadly uplifted area. Six to 10 cores were collected per site using a portable Pomeroy EZ Core Drill. A Brunton compass and inclinometer were used to determine the orientation of the beds, and the azimuth and plunge of the cores.


Cores were brought back to the University of Michigan and cut with a dual-bladed saw to 2.2 cm specimen lengths. Alumina cement was used to glue broken specimens back together. All specimens were labeled using Velvet Underglaze nonmagnetic temperature-resistant paint.

A three-axis 2G superconducting magnetometer was used to measure remanent magnetizations in a magnetically shielded room with a rest field of <200 nT. A trial run with both alternating field (AF) and thermal demagnetization showed similar magnetic directions (Fig. 2). However, the results with AF demagnetization revealed much smoother decay of the magnetization for all of the SR sites, which resulted in lower maximum angular deviations for AF-treated samples as compared to thermally demagnetized samples. Therefore, after the trial run, AF demagnetization was used to demagnetize the rest of the SR samples, with the exception of HS3 and HS4, which were thermally demagnetized. In order to reduce the acquisition of a viscous magnetization, specimens were measured directly after each demagnetization step.

A Lowrie test was performed using an ASC Scientific Impulse Magnetizer to impart a magnetization on a sample that was then thermally demagnetized. Magnetization was also applied and then demagnetized with AFs for a partial anhysteretic remanence experiment in order to examine the magnetic domain state. Magnetic hysteresis loops were acquired using a Princeton Measurements vibrating sample magnetometer at the Institute for Rock Magnetism in order to determine the distribution of magnetic coercivities. An AGICO MFK1-FA Susceptibility Bridge with CS4 furnace was utilized to determine the Curie point of the remanence-carrying magnetic mineral. Low-temperature measurements were also conducted at Institute for Rock Magnetism using a Quantum Design Magnetic Properties Measurement System to identify magnetic mineralogy by observing magnetic transitions and susceptibility dependence over a temperature range of 20–300 K.

Paleomac software by Cogné (2003) was used to analyze final demagnetization results with principal component analysis (Kirschvink, 1980). Site means were calculated by averaging the sample set directions (Fisher, 1953). In order to determine the relative timing of magnetization acquisition (pre-, syn-, or postfolding), the fold test proportionally untilts the fold limbs from measured bedding dips to horizontal (Tauxe and Watson, 1994; Watson and Enkin, 1993).


This study presents the results of 13 out of 17 sites collected from the Madison Group; the remaining four sites could not be analyzed due to spurious decay. Samples were not considered for further site analysis if the maximum angular deviation angle was greater than 20°. The magnetic characteristic directions of sites from the Mississippian Madison Group (HS3–HS4 and SR1–SR15) are shown in Table 1.

AF demagnetization of the SR sites revealed two vectors, a likely viscous present-day field component and the characteristic remanent magnetization (ChRM). The present-day field component was eliminated by 8–15 mT, and the characteristic remanent magnetization in the samples was nearly completely eliminated by 90–130 mT (Fig. 3). In the few trial samples that were thermally demagnetized, the ChRM was unblocked by 400–420 °C. The samples displayed laboratory heating-induced growth of a new mineral (suggested by a spike in the magnetic intensity) after heating the samples past 420 °C (Fig. 4B). For HS3 and HS4, thermal demagnetization was used to acquire the ChRM, which, in contrast to the SR sites, had normal polarity. In order to determine the characteristic direction, a steeper northward and normal-polarity component was first removed, as seen in Figure 3, for the HS and SR sites.

A three-dimensional isothermal remanent magnetization was applied using an ASC Scientific Impulse Magnetizer and then demagnetized, revealing goethite by the occurrence of a highly coercive component and corresponding low unblocking temperature (∼120–150 °C). Continued decay of the other size fractions to ∼550 °C supported the presence of magnetite (Fig. 4D; Lowrie, 1990). Low-temperature experiments indicated goethite at low temperatures but not at room temperature, leaving magnetite as the main remanence carrier. Partial anhysteretic remanent magnetization of a few samples as well as wasp-waisted hysteresis loops revealed a larger contribution of superparamagnetic grains as compared to solely single- or multidomain grains (Fig. 5; Tauxe et al., 1996). The following hysteresis parameters indicated a combination of superparamagnetic and single-domain grains: quality factor (Qf), remanent magnetization (Mr), saturation magnetization (Ms), coercive remanent magnetization (Brh), and coercive force (Bc; Fig. 6; Table 2; Day et al., 1977; Dunlop, 2002a, 2002b; Jackson and Solheid, 2000). Brh provides an estimate for Bcr (remanent coercive force).

Sites SR2–SR14 (excluding SR6 and SR12 for spurious behavior) have upward and steep reversed polarity directions (∼–70°), with SW to SE declinations. The two sites from the Madison Group in the Helena Salient (sites HS3 and HS4) preserve directions to the NW with normal polarities and steep directions (∼+75°). The sampled sites include five local-scale folds. Three folds (A, D, and E) resulted in synfolding remagnetization directions using the Watson and Enkin fold test to proportionally untilt the beds, while the other two (folds B and C) preserve prefolding remagnetizations (Fig. 7).


The frontal Rocky Mountains of NW Montana expose folded Mississippian Madison Group carbonates that allow for paleomagnetic analysis of local-scale folds. As determined with three-dimensional demagnetization and high-temperature susceptibility experiments of representative samples, magnetite is the magnetic carrier (Fig. 4). The derivative of susceptibility versus degrees Celsius shows a Curie temperature of ∼565 °C (Fig. 4C). The initial field cooling and ferrimagnetic signal after warming during field-cooled–zero-field-cooled (FC-ZFC) experiments did not show a clear Verwey transition (Fig. 8). The initial field cooling also depicted a hyperbolic positive increase in the magnetization upon cooling to 20 K due to the overwhelming contribution of the paramagnetic behavior within the samples (Fig. 8A). Therefore, the ferrimagnetic component (Verwey transition) was masked by the more prevalent diamagnetic/paramagnetic signal of the samples during initial field cooling and suppressed during subsequent warming and cooling curves.

Goethite was also present in the samples, as indicated by the Lowrie test and FC-ZFC curves. The Lowrie test displayed a highly coercive component that decayed by ∼120–150 °C (Fig. 4D; Strangway et al., 1968). The FC-ZFC curves revealed a high magnetization of the FC curve at 20 K as compared to the lower magnetization of the ZFC curve at 20 K after a room-temperature magnetic saturation, which is indicative of goethite (Fig. 8G; Guyodo et al., 2003).

Magnetite occurred as single-domain grains accompanied by a large superparamagnetic component in the samples from this study. Partial anhysteretic magnetization of a few samples indicated a magnetic carrier size dependence in the plot that could represent either a multidomain or superparamagnetic/single-domain component. Size dependence was revealed by a peak in the normalized magnetization of each sample in the AF window 0–50 mT (Fig. 5C). Solely stable single-domain magnetic carriers will have a normalized peak at a higher AF window (∼50–100 mT; Jackson et al., 1988). The peak in normalized magnetization that suggests a superparamagnetic/single-domain or multidomain magnetic carrier demonstrates similar behavior to what is observed in partial anhysteretic plots from the remagnetized Trenton Limestone in the eastern United States (McCabe et al., 1985).

Wasp-waisted hysteresis plots and results displayed in a Day diagram suggested a large superparamagnetic component in the samples (Fig. 6). A strong frequency dependence from 20 to 300 K also indicated a broad distribution of nanoparticle sizes (Fig. 9; Jackson and Swanson-Hysell, 2012; Worm, 1998). A robust paramagnetic contribution was seen in the hyperbolic rise of the susceptibility from ∼50 to 20 K (Fig. 9A) and was removed to observe the large distribution of nanosizes (Fig. 9B). A Curie-Weiss model paramagnetic susceptibility (k[T] = c/[T – θ]) was constructed interactively and subtracted from the measured susceptibilities. Bulk susceptibility revealed a large paramagnetic contribution of the samples after being placed in liquid nitrogen for ∼25 min (Fig. 5D). This spike in susceptibility of the samples after being cooled to ∼70 K is interpreted to be due to the paramagnetic nature of the samples responding to the presence of an ambient magnetic field when their thermal barrier was weakened (Tauxe, 2010).

The ChRM observed after removal of the present-day field component is secondary, given its steeply upward remanence (∼70°–75°) as compared to an expected shallower, North American Mississippian direction (Besse and Courtillot, 2002; Torsvik et al., 2012). Directions from the Madison Group sampled by O’Brien et al. (2007) also showed steep and upward directions (∼70°) that ranged from southeasterly to south-southwesterly (Fig. 10). An imprecise age range of Late Cretaceous to Eocene can be assigned to the remagnetization event based on comparison to expected paleomagnetic directions for North America. However, this age can be further refined by correlating published deformation ages with our reported synfolding magnetization results. The refined age of the remagnetization is Eocene (53.6 Ma) based on deformation ages in the Montana Rockies (Elliott et al., 2006; Hoffman et al., 1976; McMannis, 1965), nearby southern Alberta Rockies (Pană and van der Pluijm, 2015), and northern Wyoming Rockies (Solum and van der Pluijm, 2007).

In the study area, the remagnetization spans a reversal, as seen by down and steep directions from HS3 and HS4 and up and steep directions from the remaining sites. Samples to the east of the Rockies deformation front (site SR1) are represented by a clear present-day field magnetic component and then spurious decay of the remaining magnetization. A similar result was observed in foreland Cretaceous rocks studied by Gill et al. (2002). Thus, there was not sufficient magnetite growth to preserve Cretaceous–Eocene remagnetization away from the orogenic front in the foreland or Sweetgrass Arch.

Both syn- and prefolding remagnetizations were found in the study area by proportionally unfolding the fold limbs with the Watson and Enkin fold test method (Watson and Enkin, 1993; Fig. 7). When untilting was carried out in an asymmetric fashion, the results stayed synfolding in character. All of the fold tests were from outcrop-scale folds, which best allow an incremental fold test to determine the relative timing of the remagnetization. Folds A, D, and E (tighter folds) showed synfolding remagnetizations, whereas folds B and C (comparatively more open folds) preserved prefolding remagnetizations. With incremental unfolding in fold SR10/SR11, the kmax (maximum clustering) surpassed 100% unfolding due to a slightly more scattered site average for SR10. As seen in the stereographic projections for in situ and tilt-corrected data, the directions from each limb approach each other upon tilt correction, but they do not fully overlap (Fig. 7B). Therefore, an eigenvalue and bootstrapping fold test method (Tauxe and Watson, 1994) was employed, showing that the optimal clustering occurred shortly after 100% unfolding, at ∼104%–127% unfolding.

O’Brien et al. (2006, 2007) sampled the same Mississippian carbonates and found pre- and synfolding and tilting magnetizations in the frontal Montana Rockies. Those authors conducted a fold test when they were able to sample a fold, and where they could not sample a fold, they conducted tilt tests for samples taken from local thrusts. Similar to this study, a synfolding remagnetization was found near fold A (O’Brien et al., 2007). Within the Teton Canyon (North and South Forks of the Teton River) in the vicinity of folds B and C, O’Brien et al. (2007) found both pre- and synfolding/tilting magnetizations. A prefolding magnetization was found in the Teton anticline, i.e., the very frontal portion of the deformation belt, and a pretilting magnetization was observed in the Teton River thrusts, all comparable to the prefolding results in the same area from this study. A synfolding magnetization was also found in the Teton Canyon area, in contrast to the prefolding results from their work and this study (O’Brien et al., 2006, 2007). Tilt tests for thrusts in the Sun River Canyon (around fold D) revealed syntilting (O’Brien et al., 2006) and then pretilting magnetizations (O’Brien et al., 2007). This change was likely due to an increased number of sites that were measured and analyzed in the later paper. The pretilting result from the Sun River Canyon is in contrast to the synfolding result from fold D, likely due to the regional thrust sampling versus local fold sampling of this study.

Other studies in the area used regional tilt tests for Cretaceous-aged thrusts, as opposed to local-scale folds sampled in the Mississippian Madison Group limestone for this study. Elliott et al. (2006) found a pretilting chemical remagnetization in Cretaceous carbonate concretions from three sampled sites in the frontal Rockies near fold D from this study. However, this pretilt result may not be reliable due to the very few samples used for the tilt test and its regional application. Gill et al. (2002) also sampled Cretaceous rocks in the frontal Rockies near fold D and found pretilting or, possibly, early synfolding magnetization in a single regional fold test.

Fold test results in the study area show that folding coincided with remagnetization, whereas in the very frontal portion of the belt, it postdated remagnetization. Notably, the difference between these behaviors correlates with fold style and location in the fold-and-thrust belt. Where folding postdates remagnetization, folds are open and located in the most frontal part of the belt, whereas folds formed during remagnetization are well developed, tight, and to the west of the frontalmost portion of the deformation belt. This pattern indicates progressively younger and less-intense folding from the west to the east, a common observation in foreland fold-and-thrust belts.

Our interpretation is also supported by higher to lower ChRM magnetic intensity results from more interior locations to the most frontal locations of the sampled folds. The intensity values were determined after removal of the present-day field component to reflect the ChRM within the samples. Folds A and D are more westerly, followed by fold E, and, lastly, by folds B and C. Measured ChRM intensities are systematically lower toward the east (Fig. 11). Folds A, D, and E show higher intensities, while folds B and C have comparatively lower intensities. Site SR1, away from the deformation front, has intensities too low for an ancient magnetization to be observed, extending the pattern of lower/negligible intensities to the east in the foreland. The ChRM intensity values from site SR4 are lower than the other synfolding sites, and there is no obvious explanation as to why this occurs. However, the average ChRM intensity of folds SR3, SR4, and SR5 has a relatively high average compared with the other synfolding sites. In addition, a stronger ferrimagnetic component is depicted in hysteresis loops from folds A, D, and E as compared to samples from folds B and C, where the diamagnetic signal within the samples is more prevalent (Fig. 11).

According to a study by Hoffman and Hower (1979), maximum burial temperatures in the area were roughly 100–200 °C. Therefore, a thermoviscous remagnetization is not considered as the remagnetization mechanism based on the time-temperature relationship of magnetite (Pullaiah et al., 1975). Instead, the mechanism is interpreted as a chemical remagnetization due to growth of magnetite. Other paleomagnetic studies in the Montana Rockies have proposed illitization and fluid flow (± hydrocarbon migration) for the origin of the remagnetization in the area. In the study by Gill et al. (2002), illitization of smectite was proposed as the mechanism of remagnetization, based on the correlation of remagnetization in Cretaceous rocks in the frontal Rockies with higher percentages of illite, and in contrast, a lack of remagnetization in the foreland (Sweetgrass Arch), which is associated with higher ratios of smectite to illite. Elliott et al. (2006) also favored illitization as the mechanism for remagnetization in the area, but they were unable to constrain the paleomagnetic pole age well enough to conclude whether the characteristic magnetization was acquired prior to or during illitization. O’Brien et al. (2006, 2007) proposed remagnetization by fluids (including hydrocarbon migration) in the Mississippian Madison Group, supported by higher 87Sr/86Sr values of remagnetized carbonates and the presence of hydrocarbons in vugs of the Madison Group limestone. We conclude that rock-fluid interaction and/or an illitization-induced chemical reaction each or both may have been responsible for the observed remagnetization, with neither mechanism uniquely constrained by our study and prior work (Gill et al., 2002; O’Brien et al., 2007).

Whereas a single mechanism for the chemical remagnetization experienced by Mississippian Madison Group limestone may not be discernible, there is clearly an association with deformation-associated fluid activity. Remagnetization is a global occurrence, with individual remagnetization events typically correlated in time with local deformation events (e.g., McCabe et al., 1989; Stamatakos and Hirt, 1994; Weil et al., 2000; Van der Voo and Torsvik, 2012). In central Mexico, the Cretaceous Tamaulipas Formation displayed two distinct remagnetization events (ca. 77 Ma and ca. 44 Ma; Fitz-Diaz et al., 2014) that are correlated with a progression of deformation and fluid activity from the west to the east in the Sierra Madre Oriental (Nemkin et al., 2015). Elsewhere in the North American Cordillera, including Wyoming and Canada, remagnetization events are also Cretaceous to Eocene in age, and associated with deformation in the area (McWhinnie et al., 1990; Enkin et al., 2000; Cioppa et al., 2004).

In our study area, there is strong evidence that remagnetization progressed from west to east, weakening toward the foreland. This is recorded by magnetic intensity measurements and their diminishing values from more westerly folds in the study area to a broadly uplifted Madison carbonate sequence in the foreland. Folds A, D, and E show higher intensities (more magnetite growth), while folds B and C are comparatively lower (less magnetite growth). Site SR1, away from the deformation front, has intensities too low for ancient magnetization to be preserved, reflecting little to no observed magnetite growth (Fig. 11).

Progression of magnetite growth was likely due to orogenic fluid activity in the Montana Rockies. In northern Spain’s Cantabrian arc, Weil and Van der Voo (2002) found evidence to support a tectonically related, fluid-mediated growth of magnetite, which led to pervasive chemical remagnetizations. Another example of fluid-mediated magnetite growth is the remagnetization detected in the Devonian Swan Hills Formation of Alberta, Canada (Gillen et al., 1999). Fluid activity/presence within the North American Cordillera is supported by strontium isotope data (87Sr/86Sr), fluid inclusion analysis, and the presence of calcite veins (Lerman, 1994; O’Brien et al., 2007). We conclude that fluid-mediated growth of magnetite occurred during the Eocene, based on the correlation of synfolding remagnetization and previously determined deformation ages in the frontal region of the North American Rockies. The trend of higher to lower magnetic intensities (i.e., more to less magnetite growth) from the westerly folds to the foreland suggests that fluid activity was spatially limited, in contrast to widespread foreland activity in the Appalachians of eastern North America (Stamatakos et al., 1996; Cederquist et al., 2006; Hnat et al., 2009). Remagnetized rocks are observed in the Montana Rockies fold-and-thrust belt, but no remagnetization was detected in the foreland, reflecting the limited extent of fluid activity, mineralization, and chemical remagnetization of the Madison Group limestone.


In the study area, a pervasive Eocene remagnetization is preserved in the deformed Mississippian Madison Group limestone that constrains the structural and magnetization history of the Montana Rocky Mountains. Remagnetization records both syn- and prefolding acquisition, and a pattern of progressively less growth of magnetite toward the foreland. Prefolding magnetizations indicate a late phase of folding in the most easterly portion of the Montana Rockies, as compared to synfolding remagnetizations to the west or more interior of the fold-thrust belt.

As shown, magnetite growth resulted in a combination of superparamagnetic and single-domain grains, with a stronger diamagnetic/paramagnetic signal masking the ferrimagnetic signal in the prefolding sites as compared to synfolding sites. Given the low metamorphic conditions and likely chemical remagnetization, a local fluid origin is favored over a thermal or mechanical origin for this event. This fluid acted in a relatively narrow region, as seen by decreasing magnetic intensities toward the frontal portion of the belt and almost negligible intensities in the foreland. Therefore, we conclude that the Eocene remagnetization observed in the Montana Rockies was the result of chemically induced magnetite growth from fluids that were active during major deformation in the orogenic front, with a trend of decreasing magnetite growth toward the foreland.

The study was partially supported by the National Science Foundation, grants EAR-0909288 (Van der Voo) and EAR-1118704 (van der Pluijm), and the Turner Fund at the University of Michigan (Nemkin). Part of this work was performed as a visiting fellow at the Institute for Rock Magnetism (IRM) at the University of Minnesota. The IRM is a U.S. National Multiuser Facility supported through the Instrumentation and Facilities program of the National Science Foundation and the University of Minnesota. Special thanks go to Mike Jackson at the IRM for help with instrumentation and interpretation of the results. We thank Peter Cook for assistance with field sampling and Emily Schottenfels for help with sample demagnetization. We are grateful to the journal reviewers and Editor Weil for thoughtful comments that helped improve the paper.