Precambrian Research 278 (2016) 337–361

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Precambrian Research

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Making it and breaking it in the Midwest: Continental assembly and rifting from modeling of EarthScope magnetotelluric data

Paul A. Bedrosian

United States Geological Survey, Denver, CO, USA

article info abstract

Article history: A three-dimensional lithospheric-scale resistivity model of the North American mid-continent has been Received 23 September 2015 estimated based upon EarthScope magnetotelluric data. Details of the resistivity model are discussed in Revised 2 March 2016 relation to lithospheric sutures, defined primarily from aeromagnetic and geochronologic data, which Accepted 19 March 2016 record the southward growth of the Laurentian margin in the Proterozoic. The resistivity signature of Available online 25 March 2016 the 1.1 Ga Mid-continent Rift System is examined in detail, in particular as relates to rift geometry, extent, and segmentation. An unrecognized expanse of (concealed) Proterozoic deltaic deposits in Keywords: Kansas is identified and speculated to result from axial drainage along the southwest rift arm akin to Midcontinent rift the Rio Grande delta which drains multiple rift basins. A prominent conductor traces out Cambrian rifting Magnetotellurics Paleoproterozoic in Arkansas, Missouri, Tennessee, and Kentucky; this linear conductor has not been imaged before and Precambrian suggests that the Cambrian rift system may have been more extensive than previously thought. The high- Lithosphere est conductivity within the mid-continent is imaged in Minnesota, Michigan, and Wisconsin where it is Suture zone coincident with Paleoproterozoic metasedimentary rocks. The high conductivity is attributed to metallic sulfides, and in some cases, graphite. The former is a potential source of sulfur for multiple mineral deposits types, occurrences of which are found throughout the region. Finally, the imprint left within the mantle following the 1.1 Ga rifting event is examined. Variations in lithospheric mantle conductivity are observed and are interpreted to reflect variations in water content (depleted versus metasomatized mantle) imprinted upon the mantle by the Keweenawan mantle plume. Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creative- commons.org/licenses/by-nc-nd/4.0/).

1. Introduction scales (e.g. Frederiksen et al., 2013a; Shen et al., 2013) while detailed seismic reflection studies greatly advanced the under- The North American mid-continent presents a window into cra- standing of particular structures, such as the Mid-continent Rift ton growth and stabilization as well as the 1.1 Ga rifting event that System (MRS; Cannon et al., 1989; Woelk and Hinze, 1991; Zhu nearly tore Laurentia apart. Unique to this region is the preserva- and Brown, 1986). tion of this tectonic collage, largely unmodified by subsequent tec- Electrical resistivity is a physical property that until recently tonic events, which permits examination of how such events are has been applied only locally within the mid-continent, typically preserved in the continental lithosphere. along profiles of less than 100 km in length. EarthScope magne- Phanerozoic sedimentary rocks cover most of the North Ameri- totelluric (MT) data present the first opportunity to examine resis- can mid-continent, limiting our ability to directly sample the tivity in three-dimensions and at spatial scales necessary to underlying Precambrian collage. As such, current understanding investigate terrane-scale structure. I present here a three- of the structural evolution of the region is skewed heavily toward dimensional (3D) lithospheric-scale resistivity model derived from the northern United States and Canada, where outcrops are most EarthScope MT data collected from 2010 to 2014. I discuss details abundant. Further south, sparse borehole data provide isolated of the resistivity model in relation to lithospheric sutures that windows into the petrology, geochronology, and isotopic signature record the southward growth of the Laurentian margin in the of the Precambrian basement (e.g. Anderson, 1983; Bickford et al., Proterozoic. I examine in detail the resistivity signature of the 2015; Van Schmus et al., 2007). Geophysics plays a critical role in 1.1 Ga MRS as relates to rift geometry, extent and segmentation. these areas, and potential-field data have revealed the basic struc- The model also reveals the distribution of highly conductive tural evolution of the region (NICE working group, 2007). Recent Paleoproterozoic metasedimentary rocks in Minnesota (MN), seismic studies have examined the mid-continent at regional Wisconsin (WI), and Michigan (MI), and I discuss the origin of http://dx.doi.org/10.1016/j.precamres.2016.03.009 0301-9268/Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/). 338 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 the high conductivity. Finally, I consider what, if any, imprint has southward subduction and ultimate dextral transpression during been left within the mantle lithosphere in response to the 1.1 Ga continental collision at about 2.69 Ga. rifting event. The Superior craton began to undergo extension at 2.45 Ga In what follows I use the terms resistivity and conductivity, (Heaman, 1997) and again from 2.21 to 2.1 Ga with emplacement mathematical inverses of one another, interchangeably to describe of sills and dike swarms in Ontario and northern MN (Buchan et al., the electrical properties of the subsurface. I also use both fre- 1989; Southwick and Day, 1983). Near the southern edge of the quency and period, also inverses of one another, to describe the craton, sedimentary rocks (Chocolay and Mille Lacs Groups) were bandwidth of MT data in relation to depth of investigation. Finally, deposited unconformably on the craton during the latter exten- for compactness in describing geographic regions, I refer to states sional episode. A subsequent extensional episode, underway by by their postal abbreviation, also shown in Fig. 1. 1.91 Ga (Ojakangas et al., 2001), ended with crustal separation and creation of new ocean floor. Deposition during this time 2. Geologic and tectonic framework (Menominee and North Range Groups) began over a broad shelf along the Superior passive margin and ended within a series of The North American mid-continent region was a locus of conti- backarc basins (Animikie and Baraga Groups) as south-directed nental grown for over 1 billion years starting in the late Archean. subduction beneath arc rocks of the Wisconsin magmatic terranes The 1200+ km long Great Lakes tectonic zone (GLTZ, Fig. 1)isa (WMT) closed the short-lived ocean (Schulz and Cannon, 2007). paleosuture separating greenstone-granite terranes of the south- These sedimentary rocks host economically important banded iron ern Superior craton (2.65 Ga) from early to late Archean gneiss formations in MI, WI, MN, and Ontario. During the Penokean oro- (3.5–2.6 Ga) of the Minnesota river valley terrane (Sims, 1980). geny (1.85 Ga), calc-alkaline volcanic and intrusive rocks of the Sims et al. (1993) interpret the GLTZ to be the remnant of WMT accreted to the Superior margin along the Niagara fault zone.

Fig. 1. Precambrian geology of the mid-continent region. Underlying hillshade is the total magnetic field anomaly. Yellow and blue shaded regions denote the known extent of Precambrian clastic and igneous rocks, respectively, associated with the 1.1 Ga Mid-continent Rift System (MRS). The MRS system can be separated into the Kansas rift segment (1), the Iowa horst (2), the St. Croix horst (3), the Superior graben (4), and the Michigan rift arm (5). Red shaded region denotes the recognized extent of 1.85 Ga Penokean deformation. Regions south of the yellow line are concealed beneath Phanerozoic cover. Green shaded region denotes the extent of 1.5–1.34 Ga anorogenic magmatism that defines the granite-rhyolite province (Bickford et al., 2015). Interpreted terrane boundaries (dotted white lines) are modified after the NICE working group (2007) and Whitmeyer and Karlstrom (2007). Dashed white line is the isotopic Nd line from (Bickford et al., 2015) and dash-dot line denotes the Grenville deformation front. Abbreviated terranes or sub-provinces include the Minnesota River Valley terrane (MRVT), the Becker embayment (BE), the Wisconsin magmatic terranes (WMT) and the Marshfield terrane (MT). Named terrane boundaries include the Great Lakes tectonic zone (GLTZ), the Spirit Lake tectonic zone (SLTZ), the Niagara fault zone (NFZ), and the Eau Pleine shear zone (EPSZ). Faults associated with the MRS include the Northern boundary structural zone (NBSZ), the Thurman Redfield structural zone (TRSZ), the Belle Plaine shear zone (BPSZ), the Douglas fault (DF), the Hastings-Lake Owen fault (H-LOF), the Keweenaw fault (KF), the Thiel fault (TF) and the Isle Royale fault (IRF). Nemaha uplift (NU). Red and yellow circles indicate magnetotelluric stations shown in Fig. 3a and b, respectively. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 339

This collision also marked a transition to north-directed subduc- 1983; Hutchinson et al., 1990; Cannon and Hinze, 1992). The vol- tion along the southern edge of the WMT. Closure along this mar- ume, timing, and isotopic signature of MRS volcanic rocks are con- gin was complete by 1.83 Ga with accretion of the exotic (Archean) sistent with a short-lived thermal plume that either dissipated or Marshfield terrane and thrusting of the Paleoproterozoic sedimen- moved (Hutchinson et al., 1990; Nicholson and Shirey, 1990). tary rocks northward onto the edge of the Superior craton (Schulz Others have argued, however, that the MRS is not due to a mantle and Cannon, 2007; Ojakangas et al., 2001). The Marquette Range plume, but rather the result of passive rifting (e.g. Chase and Supergroup, bounded to the south by the Niagara fault zone, Gilmer, 1973; Gordon and Hempton, 1986), such as an oblique records metamorphism to 15+ km depth (Attoh and Klasner, continental collision along the Grenville suture and associated 1989). The preserved extent of the Penokean orogen is likely a frag- rotation of the corner of Laurentia. More recently, Stein et al. ment of what it once was. (2014) argue that the MRS is the remnant of a successful rifting From 1.8 to 1.75 Ga (geon 17; Van Schmus et al., 2007), the event between Laurentia and Amazonia, with extension along the Laurentian margin again grew southward with the addition of MRS ceasing in response to seafloor spreading between Laurentia the Yavapai terrane along the Spirit Lake tectonic zone (SLTZ). and Amazonia. The SLTZ is identified from geophysical data in northwest Iowa Thrusting from 1.08 to 1.06 Ga reactivated many of the graben- (IA) and southeast MN (NICE working group, 2007) and can be bounding rift faults, resulting in inversion of the graben, placing traced further east through central WI based on isotopic age data dense volcanic rocks higher in the section and thus giving rise to (Van Schmus et al., 2007). The composite margin experienced a prominent gravity and magnetic highs. This compressional episode widespread hydrothermal and magmatic event at 1.786 Ga is believed to be a consequence of the Grenville orogeny which (Vallini et al., 2007). Uplift and exposure of Archean gneissic rocks through far-field effects closed the rift by as much as 30 km, pre- in upper MI also occurred during this time interval and may be sumably along recently weakened zones beneath the rift grabens related to Yavapai orogenesis (Schulz and Cannon, 2007). The (Cannon, 1994). The rift is at its widest in the eastern Lake Superior accretionary growth of the Laurentian margin culminated with basin, where the alignment of the rift with respect to the Grenville the geon 16 Mazatzal orogeny (1.7–1.6 Ga). Accretion of one or front was unfavorable for compression and thought to result in more arc terranes during this period again produced a thermal strike slip motion rather than thrusting (Cannon, 1994). and deformational overprint on rocks as far north as the Niagara Excluding Cambrian rifting along the southern and eastern mar- fault zone. Within the mid-continent region, there are few con- gins of the study area (Stark, 1997), the 1.1 Ga rifting event is the straints on the location of the boundary between Yavapai and last tectonic convulsion to affect the area. Today, much of the Mazatzal rocks. MRS is covered by a thick blanket of Phanerozoic sedimentary Widespread anorogenic magmatism (e.g. Wolf River batholith rocks (Fig. 1), but observations from seismic and potential field in central WI) occurred from 1.5 to 1.34 Ga as part of a suite (the data reveal the MRS to be segmented into a series of grabens with granite-rhyolite province) that transects much of the central and the deepest rift fill (volcanic and sedimentary) preserved within southwestern United States (Anderson, 1983). Within the mid- the Lake Superior graben. The main southwest rift arm can be continent, this Mesoproterozoic magmatism spreads across a sharp traced southward into Kansas (KS), although potential field data Nd isotopic boundary between older Paleoproterozoic crust to the and limited age dating indicate the presence of 1.1 Ga mafic rocks, northwest and younger juvenile crust to the southeast (Bickford potentially related to the MRS, as far south as Oklahoma and west et al., 2015). In their interpretation, the granite-rhyolite province Texas (Adams and Keller, 1996; Keller et al., 1989). is a region of trans-Laurentian magmatism that stitches the coher- Comparatively little is known regarding the southeast, or ent lithospheric plate to the north (geon 16 and older rocks) to Michigan, rift arm which is concealed beneath the Michigan basin. juvenile oceanic crust to the south. A 250 My gap exists between South of MI, definitive age dates are absent, though the rift system the end of granite-rhyolite magmatism and subsequent tectonism, has been postulated to extend as far as Alabama (Steltenpohl et al., a time period during which the lithosphere is presumed to have 2013) along the Fort Wayne rift in Indiana/Ohio (IN/OH) and the stabilized. East Continent gravity high in Kentucky/Tennessee (KY/TN) The 1.1 Ga Mid-continent Rift System (MRS) was an abrupt rift- (Keller et al., 1982). Throughout the survey area, Phanerozoic sed- ing event spread along a 2000 km arc and cross-cutting the pre- iments form a series of basins and domes that rest upon the existing tectonic collage. It is speculated that inherited structures Precambrian collage. Structural highs include the Transcontinental played some role in focusing rifting; examples include the appar- Arch in MN, the Wisconsin dome and the Ozark dome in Missouri ent wrapping of the MRS around the expansive Wolf River batho- (MO); structural lows include the Michigan and Illinois basins. lith (Allen and Hinze, 1992) and an offset in the rift basins near the MN/IA border corresponding to the edge of the Archean Min- nesota river valley promontory (NICE working group, 2007). Rifting 3. Geophysical expression lasted for 15 My, however much of the extruded basalt was pro- duced over a 3–5 My period and was subsequently covered and The most prominent geophysical signature within the mid- flanked by a thick clastic rift-fill sequence (Davis and Paces, continent is a reflection of the MRS, seen most clearly in gravity 1990). Volcanic rocks are thickest within the Superior basin, where and magnetic field maps of the regions (Fig. 2). Thiel (1956) ini- a section as much as 30 km thick is covered by up to 12 km of sed- tially suggested that a thick volcanic and sedimentary pile was imentary rocks (Behrendt et al., 1988). The total volume of volcanic emplaced in a syncline and subsequently thrust upward along rocks has been estimated as 1.3 106 km3 (Hutchinson et al., faults to form a horst. The thick volcanic pile and segmented nature 1990) and is comparable to that of other large igneous provinces. of the rift is reflected in magnetic and gravity highs that can be Intrusive rocks are inferred to comprise much of the lower crust traced along the length of the rift. Gravity lows and regions of sub- beneath and adjacent to the rift, although surface exposure is lim- dued magnetic signature flank the central gravity high and are ited to the Superior graben (Duluth Complex, Mellen Intrusive caused by wedge-shaped basins filled with sedimentary rocks shed Complex, Nipigon embayment, and carbonatite-alkaline com- from the rift (Chandler et al., 1989). Seismic refraction models plexes along the north shore of Lake Superior). crossing the St. Croix horst and the Lake Superior graben confirm The MRS is most commonly viewed as a failed rift attributed to this first-order picture (Tréhu et al., 1991; Mooney et al., 1970). a thermal plume (100–200 °C elevation in asthenospheric temper- Seismic reflection studies in the Lake Superior graben (Behrendt ature) that impinged upon the Proterozoic lithosphere (Green, et al., 1988; Cannon et al., 1989; Mudrey and McGinnis, 2003), 340 ..Bdoin/PeabinRsac 7 21)337–361 (2016) 278 Research Precambrian / Bedrosian P.A.

Fig. 2. (a) Free air gravity anomaly and (b) total magnetic field anomaly within the study area. White circles denote long-period EarthScope MT stations. Black triangles indicate wideband MT stations from Wundermann, 1988. Thick lines denote seismic reflection profiles collected by COCORP (white), GLIMPCE (red) and industry (black). Terrane boundaries (dotted), MRS faults (solid), isotopic Nd line (dashed), and Grenville Front (dash-dot) are as in Fig. 1. Semi-transparent region shows extent of 1.5–1.34 Ga granite-rhyolite magmatism. Labels A, B, C, G, and H refer to GLIMPCE seismic profiles discussed in the text. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 341 along the Michigan arm (Cannon et al., 1991; Zhu and Brown, commercial 3D MT inversion codes have only recently become 1986), within the St. Croix horst (Chandler et al., 1989; Dickas practical on high power-desktop computers and distributed mem- and Mudrey, 2002), the Iowa rift (Anderson, 1997, 1992; Dickas ory clusters (e.g. Siripunvaraporn et al. (2005), Egbert and Kelbert and Mudrey, 2002), and the Kansas rift (Serpa et al., 1989, 1984; (2012), and Kelbert et al. (2014). Static-shift distortion, in which Woelk and Hinze, 1991) illuminate the rift system as a series of small-scale, near-surface conductivity structures distort measured broad basins bounded by prominent faults and filled with flood MT data, remains a problem in 3D, and is exacerbated by the broad basalts and overlying sedimentary rocks. Thickened crust is station spacing and limited period range of the EarthScope MT observed at some locations and is considered to reflect some com- data. 3D MT inversion is also challenging to implement in cases bination of basaltic underplating and a heavily intruded lower of non-uniform and limited site distributions, where different parts crust. of the model domain are constrained by measured data to varying Seismic reflection sections in combination with potential field degrees. In the case of the EarthScope MT data, this complexity is modeling indicate that the flanking basins contain 5–10 km of sed- limited to the edges of the array and to regions lacking data within imentary rocks; in addition, several kilometers of sedimentary Lakes Superior and Michigan. rocks overlie the volcanic pile, distributed in sag basins such as The goal of magnetotelluric inversion is to estimate the earth’s the Twin Cities basin in MN, the Ashland syncline in MN/WI, and resistivity structure from surface electric (E) and magnetic (H) field the Stratford basin in IA (Chandler et al., 1989; Anderson, 1997). measurements. The measured fields are typically transformed into Such modeling further suggests the MRS rift basins are sometimes the frequency domain and used to estimate transfer functions: asymmetric, as evidenced by rift perpendicular thickening of the  E Z Z H volcanic pile and modeled mafic intrusive roots that are offset from X ¼ XX XY X the rift axis. EY ZYX ZYY HY Prior MT studies within the mid-continent are limited. The most  extensive non-commercial study is a wideband MT profile crossing TZX HZ ¼ ½HX HY the St. Croix horst by Wundermann (1988). Modeling along this TZY profile (Fig. 2) reveals a resistive, asymmetric horst block to approximately 30 km depth and largely confined to the extent of Here, the elements Zij form the impedance tensor, a complex, the pronounced gravity high. A short MT profile crossing the frequency-dependent tensor describing the amplitude and phase MRS-related Keweenaw fault by Young and Repasky (1986) simi- relations between E and H in the earth. The complex and larly imaged a thick resistive volcanic section to the northwest of frequency-dependent vector T is known as the vertical magnetic- the Keweenaw fault within the Lake Superior graben. Up to 2 km field transfer function, or tipper, as it reflects a tipping of the mag- of conductive rift clastic rocks (40 X m) were modeled southeast netic field out of the horizontal plane due to complex geologic of the Keweenaw fault. structure. The symmetry of Z is related to the complexity of the To date, only one other study based upon the EarthScope MT subsurface resistivity distribution. In 2D, resistivity structures are data considered herein has been published (Yang et al., 2015). assumed to extend infinitely along a well-defined strike direction. Their study focuses on the sutures that stitch together the Archean This assumption, and hence 2D inversion, is applicable at a range of and Proterozoic tectonic collage and to a lesser extent on the struc- spatial scales and geologic settings, however applying it to 3D data ture of the MRS. They image prominent conductive anomalies is fraught with peril. In the case of 3D earth structure the impe- extending into the mantle along both the Niagara fault and the dance tensor is fully occupied and simplifying assumptions made Spirit Lake tectonic zones, which they interpret to dip to the south for 2D inversion are invalid. The scale of the Earthscope MT array, and north, respectively. They further interpret the lithosphere- crossing numerous tectonic elements and with no single geologic asthenosphere boundary (LAB) to be near 200 km depth, in agree- strike direction, demands a 3D inversion approach. ment with seismic LAB estimates. EarthScope seismic studies reveal a Moho that is on average 4.1. Data and errors 5 km thicker beneath the MRS along much of the southwest rift arm (Shen et al., 2013). These results are consistent with seismic The data set consists of long-period (101–104 Hz) MT data reflection studies within the Lake Superior graben (Tréhu et al., collected as part of the EarthScope USArray component since 1991; Behrendt et al., 1990). The EarthScope studies note a distinc- 2010 (Schultz et al., 2006–2018). As of December, 2014, the ‘mid- tion between the southern rift (IA and KS) where a sharp Moho is west’ footprint consists of 305 stations with nominal 70 km station observed and the northern rift (MN, WI, MI) where a more gradual spacing (Fig. 2). Data were collected at a sampling rate of 1 Hz velocity transition is observed. The latter is taken to reflect mag- using a fluxgate magnetometer and non-polarizable lead–lead matic underplating or intrusion. chloride electrodes. Sensors were oriented in geomagnetic coordi- Looking beyond the MRS, a distinct change in shear-wave veloc- nates and data were recorded for a minimum of 3 weeks; the ity is observed in northwest IA across the Spirit Lake tectonic zone resulting MT and vertical magnetic-field transfer functions are at (Shen et al., 2013). In MN, the Great Lakes tectonic zone appears as most sites well defined to 104 Hz. The MT transfer-functions were both a magnetic boundary (NICE working group, 2007) and a downloaded from the IRIS Data Management Center (www.iris. lower-crustal change in shear-wave velocity (Vs) (Shen et al., edu/spud/emtf). Time-series analysis was performed by the 2013). In WI/MI, the GLTZ is broadly coincident with a prominent National Geoelectromagnetic Facility (ngf.oregonstate.edu) using conductivity anomaly, confined to the north of the Niagara fault the multi-station remote-reference transfer-function estimation zone (Sternberg and Clay, 1977; Yang et al., 2015). program of Egbert (1997). Measured MT data, shown as apparent resistivity and phase for the principle (off-diagonal) tensor elements, are shown for several 4. Methods sites in Fig. 3. Fig. 3a illustrates the data-space variability between a pair of adjacent stations located atop the rift axis in IA and over 3D magnetotelluric inversion presents an opportunity to the nearby flanking basins. The variability at short periods suggests recover complex earth resistivity models without the assumptions lateral changes in conductivity within the upper-crustal section. and limitations inherent in two-dimensional (2D) inversion. 3D MT Fig. 3b shows two stations on either side of the Niagara fault zone. inversion, however, is computationally intensive, and non- The vastly different data at all periods indicate a strong conductiv- 342 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361

Fig. 3. MT principle impedance data at select EarthScope stations within the mid-continent study area. (a) Stations IAM36 atop the Iowa horst and IAM37 adjacent to it (red circles, Fig. 1) are significantly different at short periods, reflecting the transition from resistive Keweenawan volcanic rocks to conductive rift clastic deposits. (b) Adjacent stations MIE41 and WIF41 on either side of the Niagara fault zone (yellow circles, Fig. 1). Prominent variations at long period reflect lateral changes in conductivity across the NFZ that persist throughout the crustal column. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Fig. 4. Normalized root-mean-square (n.r.m.s.) misfit statistics for the final resistivity inverse model as a function of (a) site, (b) period, and (c) component.

ity contrast between these sites that persists throughout the mesh consists of 150 150 50 cells (1.125 million cells) with an crustal column. inner core that spans 1300 1300 km at a fixed 10 km horizontal cell size. Cell thicknesses increase logarithmically with depth from 4.2. Inversion an initial thickness of 100 m. Cells within the crustal portion of the model are on average 1.5 km thick. To satisfy boundary conditions, All inversions were carried out using the modular electromag- the vertical extent of the model is 700 km in depth while the model netic (ModEM) inversion code of Egbert and Kelbert (2012). All extends 1700 km horizontally beyond the inner core. A 100 X m 305 stations were inverted at 11 periods ranging from 8 s to homogeneous halfspace was applied as both the start and prior 17,000 s. The full impedance tensor was inverted together with model; smoothing (curvature) regularization was spatially uniform vertical magnetic field transfer functions at all sites. The 3D model in the vertical and horizontal directions and set to a = 0.4. P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 343

Specification of errors is critical in guiding the direction of any portional to the square root of resistivity and period (Orange, inversion, and can have profound effects on the resulting inversion 1989). Conservative estimates based on the longest period data model (Tietze and Ritter, 2013). In defining data errors, the relative suggest the model is sensitive to structure to at least 300 km error between tensor elements and between different periods must throughout most of the model area, the exception being in isolated be considered, as well as the effect of unknown static shifts to the regions of high lithospheric conductivity. This estimate is consis- data, which cannot be recovered during the inversion, but never- tent with a lack of visible structure in the inverse model below theless steer the inversion and distort the resulting model (Jones, 300 km, suggesting that regularization and the choice of the start 2012). Additionally, the effect of regularization, which applies con- model, rather than data sensitivity, steer the inversion in this part straints on the amount and type of structure in the resulting of the model domain. model, is intimately tied to both error specification and the result- ing data misfit. 5. Inverse resistivity model Estimated data errors, determined statistically during transfer function estimation, show a clear dependence on period (e.g. The inverse resistivity model is presented as a series of depth Fig. 3) and are distinct between diagonal and off-diagonal tensor slices through the 3D model in Fig. 5. Note the speckling apparent elements. At long periods, large error bars reflect the small number in the model, particularly within the upper crust. This speckling of transfer function estimates available for averaging, while higher reflects the physics of the MT method, and to a lesser extent the errors on the diagonal tensor elements reflect the generally smaller choice of regularization and start model. It is thus important when magnitude of these tensor elements relative to noise levels. examining the resistivity model, particularly from the surface to Although inverting measured data with their statistically- 5 km depth, not to interpret the speckled resistivity structure in determined errors would seem the most appropriate treatment, terms of geologic structure. The sensitivity of measured MT data direct application of these errors results in a coarse inverse model to regions away from the surface measurement location (both and a non-uniform data fit. This arises from limitations both in the depth and lateral distance) has a nonlinear dependence on period. data set and the ability to model it. Measured data, for example, are Long-period MT data average over a much larger volume than do sensitive to small-scale resistivity structure that cannot be mod- short-period MT data. Hence at great depth within the 3D model, eled due to mesh size limitations, cannot be constrained due to the long-period MT data are sensitive to model structure between limited data bandwidth and site spacing, and cannot be recovered the station locations. In contrast, at shallow depths, the short- due to the regularization required to solve the over-parameterized period data support volume extends only a short distance away 3D inverse problem. Furthermore, assumptions about the MT from the measurement location, while those parts of the model source fields are sometimes violated, giving rise to precise but between stations are constrained only by the starting model and inaccurate data. As a consequence of the above, an error floor is the enforced regularization. commonly applied in which data errors less than the specified The horizontal regularization can be increased to reduce speck- error floor are replaced by it. Here I have applied a 5% error floor, ling in the inverse model, however this has two undesirable effects. applied independently to each tensor element, to produce a rela- First, increased regularization results in increased smearing of tively uniform data misfit across all sites, periods, and tensor sharp lateral conductivity contrasts that the data are sensitive to. elements. Second, increased regularization, in producing a more continuous resistivity model in the upper crust, gives the false impression that 4.3. Model assessment the measured data constrain model structure between the stations. The approach taken here is to tread lightly with regard to regular- The final inverse model has a normalized root-mean-square ization, effectively highlighting regions within the 3D model con- (n.r.m.s.) statistic of 1.94, representing a 70% reduction in n.r.m.s. strained by the data, while allowing the reader to critically relative to that of the homogeneous start model. Much of the misfit evaluate, for example, whether additional structure could fall is concentrated within a small subset of sites where static shifts are between stations yet remain undetected by existing measure- visible in the data. At these sites, the n.r.m.s. for apparent resistiv- ⁄ ments. I limit my interpretation to depths greater than 1 km and ity is noticeably larger than that for impedance phase ( in Fig. 4a), to resistivity structures that are coherent over length scales of consistent with the effects of static shifts. These sites are all located 100 km or more. in northern MN and upper MI within regions of complex geology As an aid in interpreting the 3D resistivity model, laboratory where electrical conductivity is known to vary by several orders resistivity measurements were made on a suite of rock samples of magnitude over short distances (e.g. Sternberg and Clay, including Paleoproterozoic rocks of the Marquette Range Super- 1977). Higher misfits are furthermore found at shorter periods group and volcanic, intrusive, and sedimentary rocks of the MRS (Fig. 4b), and are attributed to small-scale resistivity structure that (Table 1). Details of the measurement approach are described in cannot be recovered with a 10 km cell size. Finally, the off-diagonal Bloss and Bedrosian (2015). Note that laboratory measurements impedance elements are generally fit better (n.r.m.s. of 1.56) than on hand samples may not directly reflect modeled resistivity on the diagonal impedance elements (n.r.m.s. of 2.48), as shown in the spatial scales investigated here and that the resistivity of units Fig. 4c. is in many cases based upon a single sample. As a rule of thumb, MT has a lateral resolution similar to reflec- tion seismology and a vertical resolution comparable to seismic refraction (Jones, 1987). The excellent lateral resolution arises from 6. Discussion predominantly horizontal current flow, resulting in a discontinu- ous electric field across a lateral conductivity contrast. Model res- The focus of my interpretation is on Precambrian structure and olution in MT is not a purely geometric effect (e.g., station layout evolution, however it is worth mentioning the signature of the and period range) but is strongly model-dependent. Checkerboard Phanerozoic section that blankets all but the northernmost por- tests, as employed in seismic refraction studies (Zelt and Barton, tion of the study area (Fig. 1). Examining a depth slice through 1998) are therefore of limited utility due to the strong non- the resistivity model at 1 km (Fig. 5a), the broad configuration linearity between subsurface resistivity and the amplitude of the of basins and domes is evident. The Michigan and Illinois basins surface electromagnetic fields. Depth-of-investigation is com- are all relatively conductive at the shallowest depths constrained monly expressed as a percent of electrical skin depth, which is pro- by the model (1–2 km), reflecting thick packages of conductive 344 ..Bdoin/PeabinRsac 7 21)337–361 (2016) 278 Research Precambrian / Bedrosian P.A.

Fig. 5a. Depth slices at 1 km and 4 km through the 3D resistivity model. Underlying hillshaded image is the total magnetic field anomaly. Faults associated with the MRS (solid white lines) are as in Fig. 1. Interpreted terrane boundaries (dotted and dashed white lines) are modified after the NICE working group (2007) and (Whitmeyer and Karlstrom, 2007). Black lines indicate features discussed in the text. Abbreviations include the Animikie basin (AB), Bayfield basin (BB), Duluth Complex (DC), Emerald basin (EB), Kansas basement (KB), Kansas mafic horst (KM), Lake Superior graben (LSG), mafic horst (MH), Marquette Range Supergroup (MRS), Mineola basin (MB), Stratford basin (StB), Trans-continental arch (TCA) and Twin Cities basin (TCB). ..Bdoin/PeabinRsac 7 21)337–361 (2016) 278 Research Precambrian / Bedrosian P.A.

Fig. 5b. Depth slices at 6 km and 9 km through the 3D resistivity model. Underlying hillshaded image is the total magnetic field anomaly. Faults associated with the MRS (solid white lines) are as in Fig. 1. Interpreted terrane boundaries (dotted and dashed white lines) are modified after the NICE working group (2007) and (Whitmeyer and Karlstrom, 2007). Black lines indicate features discussed in the text. Abbreviations include the Animikie basin (AB), Belle Plains resistor (BPR), Defiance basin (DFB), Duncan basin (DCB), Duluth Complex (DC), Kansas basement (KB), Kansas mafic horst (KM), Lake Superior graben (LSG), mafic horst (MH), Marquette Range Supergroup (MRS), Reelfoot rift (RR), Shenandoah basin (SB) and Wellsburg basin (WB). 345 346 ..Bdoin/PeabinRsac 7 21)337–361 (2016) 278 Research Precambrian / Bedrosian P.A.

Fig. 5c. Depth slices at 15 km and 26 km through the 3D resistivity model. Underlying hillshaded image is the total magnetic field anomaly. Faults associated with the MRS (solid white lines) are as in Fig. 1. Interpreted terrane boundaries (dotted and dashed white lines) are modified after the NICE working group (2007) and (Whitmeyer and Karlstrom, 2007). Black lines indicate features discussed in the text. Abbreviations include the Belle Plains resistor (BPR), Defiance basin (DFB), Duncan basin (DCB), Iowa intrusive complex (IIC), Lake Superior graben (LSG), mafic horst (MH), Marquette Range Supergroup (MRS), Rough Creek graben (RCG), Rome trough (RT), Shenandoah basin (SB) and Wellsburg basin (WB). P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 347

Table 1 Measured laboratory resistivity on Paleoproterozoic rocks of the Marquette Range Supergroup and on MRS rocks within or surrounding the Superior graben and adjacent to the Iowa horst. Resistivity range represents the range of resistivity values measured across frequency from 100 Hz to 100 kHz and across multiple samples if present. Colors indicate rough resistivity breakdown. Highly resistive (blue), moderately resistive (green), moderately conductive (orange), highly conductive (yellow).

Mean Era Type Unit Subunit Rock type Samples Range resisvity Bayfield Chequamegon sandstone 2 840 790-900 Bayfield Devils Island sandstone 2 2240 1540-2950 Bayfield Orienta sandstone 2 310 120-500 Oronto Freda sandstone 1 396 296-529 Oronto Nonesuch shale 2 1770 910-235 Oronto Copper Harbor conglomerate 2 1230 270-2520 interflow sandstone, Portage Lake sandstone 1 1730 1260-2120 volcanics

sedimentary sedimentary Amoco Eischeid #1, E2 sandstone 1 710 590-810 Amoco Eischeid #1, D4 siltstone 2 310 240-330 Amoco Eischeid #1, C2 black shale 2 170 150-220 Amoco Eischeid #1, C2 siltstone/shale 6 810 550-1410 Amoco Eischeid#1, C1 shale 3 1280 1110-1570

Meso-proterozoic Meso-proterozoic Portage lake altered basalt 1 820 530-1090 igneous Minong basalt 3 19900 10930- extrusive 36600 Chengwatana Volcanics basalt 1 3080 1910-4200 Duluth Complex anorthosite 1 19110 15860- 31330 igneous Duluth Complex troctolite 1 32110 26000- intrusive 39660 Mellen Intrusive Complex gabbro 1 5570 3750-6880 Baraga Michigamme pyric slate 2 0.03 0.02-0.05 Baraga Tyler argillite 1 1940 1010-2950 Animikie Biwabik ferruginous 1 4560 2370-6430 slate Menominee Ironwood ferruginous 1 8640 7634-9710 chert sedimentary sedimentary Menominee Palms argillite 1 9630 8520-13590 Paleo-proterozoic Paint River Dunn Creek pyric slate 3 0.23 0.05-0.41

sedimentary rocks deposited in these basins. In contrast, resistive near the end of rifting and during the later compressional event areas outline Precambrian structural highs with relatively thin that uplifted the horst. cover, including the Trans-continental arch in MN, the Wisconsin The shallow resistivity structure of the medial horst is all the dome, and the Ozark dome in southern MO. Given the limited more striking in comparison with the magnetic data (compare bandwidth of the EarthScope MT data, the thickness of the resistivity model and underlying hillshade in Fig. 5a). Fine scale Phanerozoic section cannot be reliably constrained by the resis- details are evident, such as faint conductors coincident with the tivity model. Stratford (StB) and Mineola (MB) basins developed atop the medial horst. The high resistivity confined to the medial horst extends 6.1. Mid-continent Rift System (MRS) well into the lower crust (Fig. 5c) and supports the presence of mafic intrusions beneath the estimated 10+ km of volcanic rocks. The 1.1 Ga MRS transformed much of the study area, extending Distinguishing the total extent of the mafic root is not possible, beyond it into Canada. The basic resistivity signature of the MRS as the high resistivity of the medial horst blends in with an other- was pointed out by Yang et al. (2015), who noted high crustal wise resistive lower crust at roughly 30 km depth. Neither at Moho resistivity over the central rift volcanic section flanked by low depths (40–50 km) nor within the upper mantle is there any dis- resistivity coincident with the flanking clastic basins. I present a tinct resistivity signature related to the Iowa horst. more detailed examination of the rift and consider its variable Gravity models suggest the flanking sedimentary basins extend expression throughout the region. I start by examining the MRS to over 10 km depth adjacent to the medial horst, and that the system in IA where its structure is most clearly evident in the southeastern Shenandoah basin (SB) extends further from the rift resistivity model, and move to progressively more complex seg- axis than the southwestern Defiance basin (DFB) (Anderson, ments, culminating in the Lake Superior graben, believed to be 1997, 1992). The flanking conductors reflect this geometry (Fig. 6a) the locus for the earliest MRS extension. I then discuss the subtle and even show a subtle divot in the otherwise conductive sedi- resistivity signature of the poorly-understood Michigan rift arm ments at the location of the Manson impact structure. The deepest and end with a discussion of intrusive rocks attributed to the MRS. basin sediments are located adjacent to the major boundary faults along both the northern (Duncan and Wellsburg basins) and south- 6.1.1. Iowa horst ern (Defiance and Shenandoah basins) ends of the Iowa horst The MRS system as imaged by electrical resistivity is most (Figs. 5b and 6). These flanking sedimentary units continue to the clearly expressed along the Iowa horst (Figs. 5 and 6). At depths southwest, where they appear to merge into a broad conductive as shallow as 1 km, conductive zones straddling a central region apron, most prominent at 6 km depth, that encircles the southern of high resistivity (>1000 X m) are evident. The latter is inter- extent of the known MRS in KS. The source of this high conductiv- preted as the volcanic section of the medial horst (MH) and possi- ity is discussed in the next section. ble intrusive rocks at greater depth. The adjacent high conductivity The 5.5 km deep Amoco Eischeid #1 well (Anderson, 1990; zones are attributed to the flanking sedimentary basins formed Anderson and McKay, 1989) was drilled into the Defiance basin 348 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361

Fig. 6. Resistivity model of the Iowa and Kansas rift basins at 6 km depth (a). Extents of the mafic horsts (black) and flanking clastic sediments (red) are taken from Chandler et al. (1989), Berendsen (1997), and Merriam (1963). Red dashed lines indicate the edge of Precambrian clastic rocks interpreted from the resistivity model. White line (dashed where inferred) denotes the down-to-the-east Humboldt fault zone that bounds the eastern margin of the Nemaha uplift. West to east profiles AA0 (b) and BB0 (c) crossing the Spirit Lake tectonic zone (SLTZ) and the Iowa horst (IH). Structural section in (c) is based upon potential field modeling along this profilebyChandler et al. (1989). Abbreviations include the Amoco Eischeid #1 well (AE1), Bayfield Group (BG), Duncan basin (DCB), Defiance basin (DFB), Kansas basement (KB), Kansas mafic horst (KM), Oronto Group (OG), mafic rocks (MR), magnetic anomaly (MA), Shenandoah basin (SB), volcanic rocks (VR) and Wellsburg basin (WB). Stratigraphic log (d) from the Amoco Eischeid #1 well in comparison to the resistivity model at the same location (AE1) and within the Manson impact structure. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) and provides a window into conductivity variations within the resistivity beneath AE1 and the Manson impact structure (Fig. 6d). clastic section. Fig. 6d presents a comparison of the 3D resistivity The current station coverage, with only 1 station located within model beneath the well (AE1) to the stratigraphic log. A general this 38-km wide Cretaceous impact, is far too coarse to discuss increase in resistivity with depth is observed, with the most the structure of the Manson Impact, aside from noting a very dif- conductive interval corresponding to the Phanerozoic section. ferent resistivity structure from that of the surrounding area. Resistivities greater than 500 X m characterize the section below 5 km depth, consistent with the intersection of gabbro (pre-rift basement) at 5.4 km in the well. In between, the Mesoproterozoic 6.1.2. Kansas rift basin clastic section has been subdivided into two intervals, believed to The structure of the Kansas rift basin is known from potential be correlative with Bayfield Group (upper) and Oronto Group field data, limited borehole data, and seismic reflection profiles (lower) sediments exposed further north (Anderson, 1997; from the Consortium for Continental Reflection Profiling (COCORP, Anderson and McKay, 1989). The resistivity model beneath the locations shown in Fig. 2). The seismic sections reveal a 3–8 km well reveals a distinction between these units, with the upper westward-deepening basin interpreted to be filled with up to Bayfield Group equivalent being 100 X m and the lower Oronto 5 km of interbedded rift volcanic rocks and clastic sedimentary Group equivalent being 500 X m. This general observation is rocks. This in turn is interpreted to be overlain by an additional noted throughout the MRS. I also note the differences in model 3 km of clastic rocks (Serpa et al., 1984). This roughly 40-km wide ..Bdoin/PeabinRsac 7 21)337–361 (2016) 278 Research Precambrian / Bedrosian P.A.

Fig. 7. Resistivity models through the northern study area. Underlying hillshaded image is the total magnetic field anomaly. White outlines indicate Paleoproterozoic supracrustal rocks (Schulz and Cannon, 2007) and include the Animikie, North Range, and Mille Lacs Groups in Minnesota and southern Ontario, and the Marquette Range Supergroup in Upper Michigan and Wisconsin. Black lines denote rift related faults, dashed where interpreted. Red circles indicate location of Terra Patrick #7-22 (TP) and St. Amour (SA) drillholes. Abbreviations include the Animikie basin (AB), Ashland syncline (AS), Bayfield basin (BB), Duluth Complex (DC), Eastern Lake Superior horst (ELSG), Emerald basin (EB), Flambeau anomaly (FA), Mellen Intrusive Complex (MC), Niagara anomaly (NA), Ottawa anomaly (OA), St. Croix horst (SCH), and the Western Lake Superior graben (WLSG). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) 349 350 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 rift basin is spatially coincident with potential field anomalies that thick blanket of clastic deposits. A modern-day analog, albeit on are interpreted to mark the central rift basin (Fig. 2). a more subdued scale, is the Rio Grande delta, which traverses The resistivity model in this region does not show a coherent, and drains multiple sub-basins of the Rio Grande rift. The maxi- highly resistive central block as imaged along the Iowa and St. mum depth of the inferred clastic sedimentary rocks is hard to con- Croix horsts. Rather, this region (KM; Figs. 5a and b, 6) is imaged strain from the MT models, however the lack of a prominent as moderately resistive from 6 to 15 km depth yet considerably gravity low, as seen on either side of the Iowa horst, suggests that less resistive than an elongated feature (KB; Figs. 5a and b, 6) clastic thicknesses in excess of several kilometers, but less than the located 60 km to the east beneath the Nemaha uplift. One explana- 10+ km modeled in Iowa by Anderson (1997) are likely. tion for the absence of a high-resistivity anomaly beneath the rift basin is the interbedding of volcanic and clastic material. This is 6.1.3. St. Croix horst in contrast to further north, where a thick volcanic and intrusive The resistivity structure of the St. Croix horst (SCH, Fig. 7) is con- pile is more characteristic of the central rift basins. Volumetric sistent with that observed along the Iowa Horst. High electrical averaging of a thick package of resistive volcanic rocks with (rela- resistivity is observed within the upper crust along the medial horst, tively) conductive interbedded clastic rocks is expected to produce in particular along the center of the horst segment, where lateral a more moderate resistivity signature than if interbedded clastic changes in resistivity coincide with the Douglas and Lake Owen- rocks were minimal or absent. Hastings faults as interpreted from seismic and potential field data The offset linear resistive high (KB) extends for roughly 200 km (Fig. 7a). High conductivity zones flank the medial horst, but in con- along the north-south trending Nemaha uplift (NU, Fig. 1), trast to the Iowa horst, appear to be more spatially limited. The straddles the Nebraska/Kansas (NE/KS) border, and is most promi- extent of the high conductivity away from the medial horst is consis- nent at upper-crustal depths (Fig. 5a and b). Interpreting this fea- tent with the extent of Keweenawan clastics interpreted from ture as rift-related mafic rock is problematic as the region is not potential-field modeling (Allen, 1994; Allen et al., 1997; Chandler anomalous in potential field maps. Furthermore, borehole data et al., 1989) and seismic refraction models (Mooney et al., 1970). constrain the extent of Keweenawan-age (Mesoproterozoic) mafic It is noted that a flanking conductor in the Bayfield basin (BB) rocks to a region west of the Nemaha uplift (Berendsen, 1997). The appears to extend 25–35 km across the Douglas fault within the distribution of Precambrian clastic rocks encountered in boreholes magnetically-defined horst block (Fig. 7a–c). This is consistent provides a clue to the origin of KB. These clastic rocks cover a large with seismic refraction studies by Mooney et al. (1970) who model area extending both west and east of the rift basin, but are not low velocities, consistent with Keweenawan sedimentary rocks, encountered in drillhole data along the length of the Nemaha east of the Douglas fault in what they interpret to be a narrow gra- uplift. The Nemaha uplift zone is interpreted to have been formed ben within the fault zone. Wideband MT data over the western in post-Keweenawan times, and Berendsen (1997) suggests that part of the SCH (site locations shown in Fig. 2) further constrain clastic rift rocks were removed following uplift and deposited in the structure in this area (Wundermann, 1988). My modeling of basins further to the east. In this interpretation, the linear resistor, this wideband data over the SCH reveals a 100–500 m thick, elec- KB, is nothing more than a basement high, with a resistivity value trically conductive layer at 1.5 km depth that pinches out to the similar to nearby Mazatzal crust, but surrounded by a thick sea of east beneath the center of the horst block. The model based upon conductive Keweenawan clastic rocks. This interpretation is sup- EarthScope MT data, with more limited bandwidth, produces a ported by drillhole depth-to-Precambrian estimates, which reveal blurred image of this conductor. Wundermann (1988) interprets an uplifted bench to the west of the Nemaha uplift with more than this conductive layer as a wedge of interflow clastics, an interpre- 600 m of offset across it (Berendsen, 1997). tation with which I concur, although the modeled resistivity (10– The conductive regions surrounding KM and KB are thus inter- 100 X m) for this layer is less than laboratory measurements on preted as Precambrian clastic material shed from the uplifted rift interflow clastic rocks (Table 1). basin, yet a distinct difference is observed in the configuration of The interpreted clastic wedge appears localized to the western these sedimentary rocks in comparison to those flanking the Iowa half of the SCH. The fine-scale texture in the magnetic-field data horst. In Kansas, the high conductivity forms a broad fan that over the SCH (Fig. 2b), in particular encircling the nearby Twin extends quite far (200+ km) from the rift basin and appears deep- Cities basin and Ashland syncline, suggests the current erosional est along an arc encircling the southern end of the Kansas rift basin level is at or near the top of the original volcanic pile. As such, a (Fig. 6a). Drillholes have encountered thick packages of conglomer- clastic wedge at 1.5 km depth may mark a time interval when rift ate, arkosic sandstone, siltstone and shale which have tentatively volcanism was waning and when sedimentation was increasing. been correlated with Oronto and Bayfield Group rocks Alternatively, it may indicate a complex rift evolution with alter- (Berendsen, 1997). Where drilled, thicknesses in excess of nating periods of volcanism and sedimentation, in contrast to the 1000 m have been encountered yet do not penetrate the full sedi- more widely held view that MRS volcanism largely preceded sedi- mentary section (Berendsen, 1997). The lack of sedimentary rocks mentation, with volumetrically minor amounts of interflow clastic along the axis of the Nemaha uplift is consistent with post- sediments. Keweenawan erosion during uplift. The spatial extent of the high The rift-flanking conductors are most prominent along the conductivity fan is consistent with limited drillhole observations southern end of the St. Croix horst, within the basins adjacent to (solid red lines; Fig. 6a), yet can be extended beyond into southeast the Belle Plaine fault (Fig. 5a and b). I speculate that this region NE (dashed red lines), encircling both the Kansas rift basin and the may have been a local depocenter for Bayfield Group sediments southern end of the Iowa horst, an area of over 80,000 km2. shed southward from the St. Croix horst and northward from the I speculate that this vast zone of high conductivity is the result Iowa horst. In contrast to the flanking basins, the Twin Cities basin of axial sediment transport along the western rift arm during late- (TCB, Fig. 5a) is of moderate resistivity (100–300 X m), distinct stage rifting or subsequent uplift of the individual rift basins. While from the resistive mafic horst it sits upon, yet less conductive than there is evidence for reverse faulting (and hence uplift) along the clastic rocks in the flanking basins. This is interpreted to reflect a Kansas rift (Woelk and Hinze, 1991), the magnitude of uplift is pre- distinction between the older Oronto Group sedimentary rocks, sumed to be significantly less than that along both the Iowa and St. typically located within the grabens and atop the horsts, and the Croix horst blocks. Thus it is plausible that the terminus of the younger Bayfield Group (and equivalent) sedimentary rocks that western rift arm remained structurally lower than the northern are volumetrically the largest component of the flanking basins reaches, and that south-directed sediment transport resulted in a (Chandler et al., 1989; Mooney et al., 1970). Laboratory resistivity P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 351 measurements on the Freda and Copper Harbor Formations, volu- and along its northern shore in Ontario. The eastern half of the metrically the largest members of the Oronto Group, were found basin in particular is poorly constrained, although stations on Isle to be on the order of 400 X m and 1500 X m, respectively (Table 1). Royale and at the tip of the Keweenaw peninsula dramatically Examination of the resistivity model in the shallow subsurface improve resolution of structural variations within the basin. A beneath the Bayfield peninsula (Fig. 7a), the type locality for the first-order variation is observed in upper-crustal resistivity Bayfield Group, reveals the high conductivity most commonly seen between the western and eastern portions of the Lake Superior flanking the southern rift basins. Nearby, the Douglas fault sepa- basin (Fig. 7a–d), with the western half being on average more rates Bayfield Group rocks to the north from Oronto Group rocks resistive (1000+ X m) than the eastern half (100 X m). The bound- to the south. Corresponding model conductivities are higher to ary between the two regions is not well resolved within the resis- the north of the fault over the Bayfield Group rocks. tivity model, but is consistent with the location of the Thiel fault, Laboratory measurements on Bayfield Group rocks an interpreted down-to-the-east fault zone hypothesized by (300–2250 X m) do not reveal any distinct conductivity difference Hinze et al. (1966) from potential-field data. Subsequent seismic between them and Oronto Group rocks. Two lines of evidence sug- imaging and potential-field modeling along GLIMPCE line G gest, however, that the Bayfield Group is on average more conduc- (location in Fig. 2) provides confirmation of this fault, which is tive that the Oronto Group. First, electrical logs (M. Humphrey, interpreted to separate a relatively thin package of sedimentary pers. comm.) from the St. Amour well (SA, Fig. 7), located east of and volcanic rock to the west from a pronounced eastward- the Keweenaw peninsula, contain over 600 m of Jacobsville Sand- thickening package of sedimentary and volcanic rock (Cannon stone, equivalent to Bayfield Group, characterized by a resistivity et al., 1989; Thomas and Teskey, 1994). of 10 X m. Second, correlation of the 3D resistivity model with I interpret the first-order resistivity transition that occurs near the AE1 well (Fig. 6d) shows higher resistivity for the lower clastic the Thiel fault to reflect a transition from the western basin where sequence (units B–D), attributed to Oronto equivalent, than with post-rift compression has resulted in significant uplift, akin to the the upper clastic sequence (units E–G), attributed to Bayfield Iowa and St. Croix horsts, to the thick eastern basin, where I spec- equivalent rocks (Anderson, 1997; Anderson and McKay, 1989). ulate that graben-bounding faults were not as favorably aligned so This association is loosely borne out by measurements on core as to accommodate reverse motion (Cannon, 1994). In this context, samples from the Eischeid well (Table 1). the low resistivity within the eastern basin reflects preservation of Somewhat enigmatic is the Ashland syncline (AS, Fig. 7a), a thick sedimentary package, whereas the high resistivity in the which, based upon seismic constraints and potential field models, western basin reflects the relatively thin veneer of sedimentary contains 2–2.5 km of Oronto Group sedimentary rocks (Chandler rocks and shallow and volcanic rocks preserved following com- et al., 1989). The Terra-Patrick #7-22 well, located within the Ash- pression. Hereafter I refer to these subdivisions as the western land syncline immediately south of the Douglas fault (TP, Fig. 7) and eastern Lake Superior graben (WLSG, ELSG). drilled over 1400 m of Oronto Group rocks and did not encounter Independent evidence for this west to east asymmetry comes igneous basement. In contrast to other medial basins (Stratford, from seismic and potential-field data. East of the Thiel fault, Mineola and Twin Cities basins), the resistivity model over the GLIMPCE line F imaged a thick, largely symmetric graben inter- Ashland syncline is resistive and virtually indistinguishable from preted to be filled with 12–14 km of clastic rocks atop a 20-km the mafic horst it sits upon. This may reflect the particular compo- thick volcanic pile (Behrendt et al., 1990, 1988). In easternmost sition of the Ashland syncline rocks, which throughout much of the Lake Superior, an additional seismic line and potential-field model- syncline is exclusively Copper Harbor Formation (a conglomerate). ing by Mariano and Hinze (1994a,b) indicates a similar structure, Given the upper value of measured resistivity for the Copper with a 150 km wide graben extending to 25 km depth. In contrast, Harbor Formation (2500 X m, Table 1), it is plausible that the Ash- in the western part of the lake, GLIMPCE line C, in particular the land syncline has little resistivity contrast with the underlying vol- northern part, images a relatively thin (5–7 km) package of sedi- canic rocks. mentary and volcanic rock interpreted to sit atop Archean base- ment (Cannon et al., 1989). 6.1.4. Lake Superior graben Within the WLSG, interpretations of GLIMPCE line C infer Of all the rift basins within the MRS, the Lake Superior graben is reverse motion upon a fault, termed the Ojibwa fault by the best characterized. Seismic reflection surveys by the Great Sexton and Henson (1994), that runs parallel to the Keweenaw Lakes International Multidisciplinary Program on Crustal Evolution fault and separates a deep and narrow graben to the southeast (GLIMPCE) and industry data (Mudrey and McGinnis, 2003) have from an Archean structural high to the northwest. The extent imaged the structure of the rift beneath the lake in considerable and importance of this fault is subject to debate, with Allen detail (profile locations are shown on Fig. 2). Together with poten- et al. (1997) interpreting the Ojibwa fault as a local feature tial field models (Allen et al., 1997; Mariano and Hinze, 1994a, bounding the southern edge of the Isle Royale anticline. The 1994b; Thomas and Teskey, 1994), these studies illuminate the resistivity model suggests the Ojibwa fault to be a regional fea- Lake Superior basin to be a complex graben, bound loosely ture, as the linear resistivity contrast can be traced for 125 km between the Keweenaw and Isle Royale fault systems, and seg- from the eastern edge of the Bayfield peninsula to just south mented along is axis by the Thiel fault. The graben, attaining a of Isle Royale (Fig. 7a–c). On the southeast side of the WLSG, maximum depth in excess of 30 km, is filled with a thick lower vol- pronounced reverse motion has occurred along the Keweenaw– canic unit and an overlying blanket of Oronto and Bayfield Group Marenisco–Pelton Creek fault system, with 30 km of vertical off- sedimentary rock. A north to south asymmetry is also evident, with set documented (Cannon et al., 1993). the northern and southern margins of the graben dominated by In greater detail, the linear resistivity boundary interpreted as plutonic and volcanic suites, respectively (Thomas and Teskey, the Ojibwa fault separates more resistive rocks to the southeast 1994). I examine details of the above model for the Lake Superior from a region of more moderate resistivity to the northwest basin through the lens of the resistivity model. In particular, I argue (Fig. 7a and b). In the upper crust, I interpret the contrast to for a structural asymmetry between the western and eastern reflect conductive Bayfield Group juxtaposed across the Ojibwa halves of the basin, as well as a horst-and-graben type segmenta- fault against more resistive Oronto Group, consistent with the tion within each half of the basin. interpretation of Thomas and Teskey (1994) who modeled a den- The resistivity model over the Lake Superior basin (Fig. 7) has sity contrast along GLIMPCE line C at the Ojibwa fault. At greater limited resolution due to a lack of MT stations both within the lake depth, the region to the northwest includes the northern Bayfield 352 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 peninsula and what have been termed White’s ridge and Grand nawan extension, with the southwest flank as much as 3 km dee- Marais ridge, regions of pronounced gravity lows (Fig. 2a) inter- per than the northeast flank. preted by Allen et al. (1997) to be structural highs at the time Resistivity imaging of the MI rift arm is hindered by in excess of of rifting. This region is bounded by the Duluth Complex to the 4 km of Phanerozoic rocks (Hinze et al., 1975) within the roughly northwest, the Ojibwa fault to the southeast, Isle Royale to the circular MI basin (Fig. 5a). At mid-crustal depths, a sinuous con- northeast, and the Douglas fault to the southwest. The resistive ductor flanked by more resistive material can be traced through WLSG is observed to broaden considerably near Isle Royale, lower MI (Fig. 5c). As such, this structural pattern is similar, albeit bound between the Keweenaw and Ojibwa faults in the south- much deeper, than the geophysical signature in the Lake Superior west, yet extending from the Isle Royale fault to the Keweenaw graben. The conductor corresponds closely in shape, but appears fault further east (Fig. 7a and b). Approaching the Thiel fault from offset in location, from the MI gravity high, the outline of which the west, a compact conductive body just north of the Keweenaw is shown in Fig. 5b and c as a solid black line. fault (Fig. 7a–c) correlates well with a southward thickening I speculate that the conductor at these depths is the Mesopro- package of Bayfield Group rocks interpreted to extend to 10+ terozoic clastic section interpreted by Zhu and Brown (1986). The km depth along GLIMPCE line A (Cannon et al., 1989; Thomas conductive ribbon is observed, however, to extend significantly and Teskey, 1994). deeper than the base of clastic rocks as interpreted from the seis- I now turn to the ELSG as imaged by the resistivity model. mic sections. This may be an artifact of the inversion, as MT data Within the ELSG, the resistivity contrast is reversed from that are only weakly sensitive to the base of a conductor and the regu- observed along the Iowa and St. Croix horsts, which are marked larized inversion approach tends to smear out resistivity gradients. by a transition from high resistivity inside the horst to more mod- It is possible, though, that the deeper volcanic section interpreted erate or even low resistivity outside the horst (Fig. 3a). The package by Zhu and Brown (1986) remains electrically conductive due to of sedimentary rocks within the ELSG is less resistive than the the interbedded clastics rocks, similar to my interpretation of the surrounding Archean crust. In this context, the southwestern KS rift basin. In this case, the conductor would represent the boundary of the ELSG is most clearly seen as a linear, northwest- roughly 12-km thick syn-rift sequence beneath the Phanerozoic trending resistivity contrast (Fig. 7a–d) that cuts across upper cover. The cause of the offset of the conductor relative to the MI Michigan near the St. Amour (SA) deep drillhole. This boundary gravity high is unknown, but may reflect the asymmetry in the rift is sub-parallel to a linear boundary in both gravity and magnetic- graben as noted from the seismic interpretation. field data (Fig. 2). The SA drillhole lies within the graben and The presence of a thick clastic sequence is at odds with penetrated over 1400 m of Bayfield and Oronto Group equivalent potential-field modeling by Hinze, Kellogg and O’Hara (1975), sedimentary rocks before encountering rhyolite at a depth of who in the southeast corner of the state modeled a volcanic section 2100 m (Ojakangas and Dickas, 2002). The northeastern boundary at the base of the Phanerozoic section with no intervening Meso- of the ELSG is less clear, but an eastward increase in resistivity is proterozoic clastic rocks. I note that their interpretation is based evident near the eastern edge of the station array, in accordance upon modeling a single profile crossing a circular magnetic high with the inferred eastern rift margin as interpreted from that is not representative of the MRS in lower MI. I thus interpret magnetic-field data (dashed lines in Fig. 7). the sinuous conductor to represent the resistivity expression of A final detail within the ELSG is the complexity in the resistivity the MRS in lower MI. The extent of the conductor cannot be traced model across the axis of the graben. Given the rift boundaries east of the Grenville Front in the resistivity model, although this is described above, the eastern portion of the graben appears more near the eastern edge of the station array. Further south, no resistive than the western part from 2 to 6 km depth. Similarly, distinct conductivity anomalies are imaged coincident with gravity gravity data suggest a west-to-east transition, with the more resis- highs over the Fort Wayne rift and the East Continent Gravity High. tive eastern zone corresponding to higher gravity. Mariano and Hinze (1994a) interpreted a fault along this boundary based on a 6.1.6. Intrusive rocks change in seismic reflectivity. I speculate that west-to-east seg- Intrusive rocks of the MRS include alkaline intrusions, dikes, mentation within the ELSG reflects a horst-and-graben structure, sills, and layered intrusive complexes. Alkaline intrusions of MRS with greater uplift occurring in the east. age, such as the sub-circular Coldwell Complex (Heaman and Machado, 1992), are only known definitively along the north shore 6.1.5. Michigan rift arm of Lake Superior in Canada. As the MT station array does not extend The eastern, or Michigan, rift arm is the least constrained rift this far north, the resistivity model does not image these intru- segment both in terms of structure and overall extent. A gravity sions. Similarly, the laterally extensive Nipigon sills in Ontario high, similar to the Iowa and St. Croix horsts in character but sub- (Heaman et al., 2007) fall outside the model domain. dued in amplitude due to the overlying Michigan basin sedimen- The layered mafic Duluth Complex in northern MN is the largest tary rocks, can be traced through lower MI up to the Grenville known intrusive complex of MRS affinity. It is characterized by Front (Fig. 2). Some authors argue that similar gravity anomalies, both a prominent gravity high and a complex magnetic field anom- including the East Continent Gravity High in KY/TN extend the aly distinct from the surrounding region (Fig. 2). In the resistivity MI rift arm further south (Keller et al., 1982), potentially as far as model, the Duluth Complex is highly resistive through the entire Alabama (Stein et al., 2014). I do not examine the southern extent crustal column (DC, Fig. 7). Its western edge is particularly well of the MI rift arm, instead limiting discussion to the rift’s expres- defined, where a strong conductivity contrast separates it from sion in lower MI. Paleoproterozoic metasedimentary rocks of the Animikie basin. A handful of wells encounter Keweenawan-age volcanic rocks Other intrusive complexes of the MRS include the Mellen Complex beneath the MI basin (Catacosinos, 1973; Fowler and Kuenzi, in northern WI and the Crystal Lake gabbro in Ontario. The Mellen 1978). Additionally, several short COCORP seismic reflection pro- complex (MC, Fig. 7) is similarly resistive, although it cannot be files (locations in Fig. 2) cross the MI rift arm and reveal two inter- clearly distinguished from adjacent MRS volcanic rocks that are preted Mesoproterozoic packages – a weakly reflective upper also resistive. clastic section (4 km thick) and a more reflective lower section Rift related intrusives beneath the concealed portions of the (8 km thick) interpreted as interbedded volcanic and clastic rocks MRS are uncertain. The Northeast Iowa intrusive complex (IIC) (Brown et al., 1982; Zhu and Brown, 1986). The authors suggest has been speculated to be of MRS age as it is situated, similar to that the MI rift graben is asymmetric due to a period of late Kewee- the Duluth Complex, adjacent to the rift axis (Anderson, 2006). P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 353

The potential-field signature of the IIC, in particular the high grav- ity anomaly (Fig. 2a), suggests a large volume of mafic to ultra- mafic rock (Drenth et al., 2015). The IIC stands out in the resistivity model, particularly in the lower crust (black dashed line; Fig. 5c) where it is imaged as a coherent highly resistive region covering much of eastern IA. While the age of the IIC is unknown, the resistive anomaly appears to span the Yavapai-Mazatzal boundary, suggesting it postdates 1.6 Ga. The areal extent of the high resistivity region centered on the IIC is more than five times as large as that of the Duluth Complex. Mafic dikes are preserved in several groupings around the periphery of the MRS, including in northeast MN, Upper MI, and central WI. Volumetrically, these intrusive rocks are a minor com- ponent and are neither expected, nor observed, to have any expres- sion in the resistivity model.

6.2. Paleoproterozoic basins

The most intense conductivity anomalies in the entire study area are attributed not to the Mesoproterozoic MRS, but to Paleoproterozoic metasedimentary rocks in MN, WI, and MI formed during the 1.85 Ga Penokean Orogeny and subsequent Fig. 8. Resistivity model at 2 km (a) and 4 km (b) depth within the northern study tectonic overprints. Where these rocks are exposed (white outlines area. Underlying hillshaded image is the total magnetic field anomaly. White in Fig. 7), patterns of stratigraphy, sedimentation, structure, and dashed line indicates the southern edge of rocks of the MRS. Gray shaded regions magmatism all point to an early extensional period that preceded indicate mapped meta-graywacke of the Michigamme Formation (Cannon and the intense compression at the end of the Penokean orogeny (e.g. Ottke, 1999). Black solid lines indicate the extent of various AEM surveys flown within the study area; white outlines denote the approximate extent of conductive Sims and Peterman, 1983). The ultimate origin of the electrically- anomalies identified in such surveys. Flambeau anomaly (FA), Ottawa anomaly conductive metasedimentary rocks is a sequence of carbonaceous (OA), Niagara anomaly (NA). marine sediments deposited in tectonic foredeeps along the south- ern edge of the Superior craton toward the end of the Penokean Orogeny (Boerner et al., 1996; Ojakangas et al., 2001; Schulz and Cannon, 2007). These foredeep basins are thought to have localized aly. The authors modeled the data using a series of vertical along the Great Lakes Tectonic Zone, a pre-existing zone of weak- conductive sheets of infinite along-strike extent; their best fit solu- ness in the Archean crust (Sims and Peterman, 1983). The sedimen- tion includes five parallel conductive sheets extending from close tary rocks were in many places deformed and the carbon to the surface to 16 km depth and with a conductivity-thickness metamorphosed to graphite during Penokean compression starting product of 100 Siemens. around 1.84 Ga and again during post-orogenic collapse in mid- Additional constraints on the geometry and extent of the FA Geon 17 times (Holm and Lux, 1996; Schneider et al., 2004). come from airborne electromagnetic (AEM) surveys flown in the In MI and WI, the degree of deformation decreases to the north, 1970 s and 1980 s (black outlines, Fig. 8). An AEM survey extending with metasedimentary rocks in the south adjacent to the Niagara from 90.5°Wto90°W, also presented in Sternberg and Clay (1977), fault zone experiencing severe deformation during the suturing revealed a series of narrow, elongate conductors striking N70°E of this region to the Wisconsin magmatic terranes at the end of and extending the width of the survey area (white outlines, the Penokean Orogeny (Cannon and Ottke, 1999; Schulz and Fig. 8). As AEM systems have limited depth-of-investigation, the Cannon, 2007). In MN, deformation is also more intense in the presence of conductive anomalies implies that they lie within a south, where a series of thrust panels give way to essentially unde- few hundred meters of the surface. Just east of 90°W, an AEM sur- formed strata within the Animikie basin. These variations in defor- vey flown over the Lac du Flambeau reservation imaged similar lin- mation have resulted in variations in thickness as well as ear, anastomosing conductive anomalies over the northern third of metamorphic grade, with Paleoproterozoic metasedimentary and the reservation; this same survey notes a distinct lack of conduc- metavolcanic rocks thickening to the south, with thickness esti- tive anomalies south of a boundary that coincides with the Niagara mates of up to 9 km in the Iron River and Menominee ranges fault (B. Smith, pers. comm.). To summarize past studies, the high (Morey, 1996; Ojakangas et al., 2001). conductivity of the FA has been documented along much of its With little exception, all the regions of extremely high electrical length and consists of a series of deep-seated, elongate, sub- conductivity (<1 X m) fall within mapped Paleoproterozoic areas. vertical conductors that come within several hundred meters of Note, however, that there are regions of mapped Paleoproterozoic the surface, and falls within a region bounded to the south by rocks that are not electrically conductive (Fig. 7). These variations, the Niagara fault. as will be discussed, are related to differences in the source rocks in The northern edge of the FA correlates loosely with the bound- these regions as well as differences in the peak metamorphic ary between the Michigamme structural terrane to the north and grade. I begin by examining the individual high conductivity zones. the Watersmeet and Park Falls structural terranes to the south (Cannon and Ottke, 1999). Primary differences between these 6.2.1. Northern Wisconsin and Upper Michigan structural terranes include the thickness of the metasedimentary The Flambeau Anomaly (FA) is an elongate electrical conductor rocks, the type of Archean rocks they rest upon (Superior granite/- in northern WI that trends west-northwest and extends for greenstone versus Minnesota River Valley gneiss) and the degree of approximately 130 km along strike (Fig. 8). The FA was originally metamorphism they have undergone. The Michigamme terrane is described by Sternberg and Clay (1977), who, using a series of largely greenschist facies whereas the Watersmeet and Park Falls long-offset DC resistivity soundings and audio-frequency magne- terranes are both amphibolite facies, containing kyanite and silli- totelluric data, modeled the western half of the conductivity anom- manite index minerals, respectively. 354 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361

Fig. 9. Resistivity model at 2 km (a), 6 km (b), 9 km (c), and 15 km (d) beneath central Minnesota. Underlying hillshaded image is the total magnetic field anomaly. White lines indicate mapped extent of Paleoproterozoic rocks. Gray shaded region denotes the extent of a fold and thrust (F&T) belt incorporating Paleoproterozoic and Archean rocks. Black line traces the Malmo Discontinuity (MD), dashed where inferred. Animikie basin anomaly (ABA), Nimrod outlier (NO), Long Prairie basin (LPB), McGrath gneiss dome (MGD), St. Croix horst (SCH), Duluth Complex (DC).

The EarthScope resistivity model is consistent with the results northwest edge of the anomaly appears to extend 10–20 km of past studies of the FA but for the first time images the full extent beneath younger rocks of the MRS (Fig. 8). A less prominent high of the anomaly, albeit on a coarse spatial scale. The FA is seen to conductivity zone is imaged beneath the east central Keweenaw extend from 91°W to 89.5°W, is 20 km wide and, within the reso- peninsula. Young and Repasky (1986) modeled a short MT profile lution of the model, lies entirely to the north of the Niagara fault. crossing the Keweenaw peninsula at this location, and similarly Examining model conductivity as a function of depth, the FA interpreted a high conductivity zone at depth as Michigamme For- appears to be subvertical. One additional observation is that the mation overlain by volcanic and clastic rocks of the MRS. Several central portion of the FA becomes somewhat less conductive start- AEM surveys have been flown along the northern, western, and ing at a depth of 6 km, whereas the west and east ends remain eastern boundaries of the OA (Ensign, 1981; Geoterrex Ltd., 1982, highly conductive (Fig. 7c). 1981; Heran and Smith, 1980). Regions of elevated near-surface In WI and MI, the 3D resistivity model images two additional conductivity identified by these surveys are outlined in white in conductivity anomalies attributed to Paleoproterozoic rocks. Fig. 8, and show a strong correlation with both the edges of the Northeast of the FA is an extensive region of elevated conductivity OA and the extent of the mapped Michigamme Formation. The that lies within an area of mapped Michigamme Formation meta- OA falls across a transition from greenschist- to amphibolite- graywacke (Fig. 8). This anomaly, which I term the Ottawa anom- facies rocks (Cannon and Ottke, 1999; Schneider et al., 2004). aly (OA), as it fall largely with the Ottawa National Forest, has an The final conductivity anomaly is an intense highly conductive approximate east-west elongation and is 125 km in length by zone spanning the WI/MI border in the vicinity of Niagara, WI. 45 km in width. The OA is bound to the north and east by Archean Herein called the Niagara anomaly (NA), the anomaly is spatially rocks of the Superior Province and Minnesota River Valley terrane, coincident with the Iron River structural terrane (Cannon and respectively, with the GLTZ forming its northeast boundary. The Ottke, 1999) and again overlaps Michigamme Formation P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 355 meta-graywackes. The NA appears to extend for 60 km along a correlative to the Michigamme Formation in MI/WI. Unlike the northwest trend and is loosely bound to the south by the Niagara Michigamme, however, the Virginia and Rove Formations are fault. Its across-strike extent is roughly 40 km. Similar to the OA, essentially undeformed except in proximity to the fold and thrust the NA overlaps both greenschist-facies rocks in the south and belts in the south (Fig. 9). amphibolite-grade rocks in the north. No independent geophysical Looking to the resistivity model in central MN, there are 2 dis- data are available to further constrain the geometry and nature of tinct regions of anomalous conductivity that are similar to those the NA, however it is similar to the FA and OA both in terms of con- imaged in MI/WI. The first is located between 2 and 6 km depth ductivity (<1 X m) and spatial scale. Morey (1996) estimates up to beneath the central Animikie basin, with the deepest conductive 9 km of stratified Paleoproterozoic rock in this area. This, as well as region in the north. The anomaly is slightly elongated in a north– the consistent overlap between the conductive anomalies, mapped south direction, covering an area that is 120 km by 90 km. Notably, Michigamme Formation rocks, and high metamorphic grades sug- this anomaly does not appear to extend east of the mapped extent gest a common source for all three anomalies. of Paleoproterozoic rocks in MN, either beneath the Duluth What is the cause of the high electrical conductivity and is it Complex or within western Lake Superior. This is consistent with consistent with the tectonic evolution of the area? Sternberg and potential-field modeling by Thomas and Teskey (1994), who termi- Clay (1977) attribute the enhanced conductivity to a series of anas- nated the dense Animikie Group section at the lakeshore. The sec- tomosing steeply-dipping graphitic ‘stringers’ that extend for tens ond conductivity anomaly is considerably deeper (9–15 km), of kilometers. Although the region is covered by glacial till, drill elongated in an east-west direction (60 km by 40 km), and located hole data intersect 1–15 m thick zones of ‘pencil-quality’ graphite, further south beneath the McGrath gneiss dome within the south- and measurements of the most conductive core intervals average ernmost structural panel of the fold and thrust belt (Fig. 9). 0.03 X m(Sternberg and Clay, 1977). Analysis of drillhole data Given the geophysical similarity between the Animikie basin within the northern portion of the Lac du Flambeau reservation anomaly (ABA) in MN and the Flambeau, Ottawa, and Niagara reveals 1–3 m thick zones of graphite in numerous holes below anomalies in MI/WI, as well as the correlative stratigraphy 30 m of glacial till (B. Cannon, pers. comm). The holes intercept- between the Virginia and Rove Formations in MN and the ing graphite are spatially coincident with AEM conductivity Michigamme Formation in MI/WI, it is tempting to attribute meta- anomalies, while those to the south of the Niagara fault neither morphic graphite to the cause of the ABA. Key differences, how- show graphite nor are located near any AEM conductivity ever, include the depth extent of these anomalies and the anomalies. metamorphic grade of the respective regions. The Animikie Group Graphite is a plausible explanation for the high conductivity within the region of the ABA is weakly metamorphosed, in contrast anomalies, provided the source rocks contain carbonaceous to amphibolite-facies rocks co-located with the anomalies in material and these rocks have undergone metamorphic conditions MI/WI. The ABA is primarily limited to the upper 4 km and is sufficient to convert this material into graphite. The formation of sub-horizontal in aspect, while the other anomalies extend to graphite from carbon-bearing sedimentary rocks is highly depen- mid-crustal depths and are consistent with vertically-dipping dent on temperature. Fully-ordered graphite is only found in meta- structure. Additionally, there is no evidence for graphite within morphic rocks of amphibolite grade or higher, where it is correlated the Animikie Group rocks. with peak metamorphic temperatures in excess of 400 °C and lar- Finally, the measured data over the ABA are distinct, for exam- gely independent of pressure (Landis, 1971). A Raman spectroscopy ple, from data measured over the Ottawa Anomaly (OA, Fig. 8). study of carbonaceous material at various metamorphic grades Fig. 10 shows phase tensor ellipses at periods corresponding to found only ordered graphite in rocks that experienced peak meta- depths of 2–4 km and at MT stations located over the center of morphic temperatures in excess of 650 °C(Beyssac et al., 2002). I the ABA and OA. Phase tensors provide a view of the ‘dimensional- thus consider 400 °C to be the temperature at which graphite begins ity’ inherent in measured MT data and by proxy the complexity of to form and 650 °C to be the temperature at which carbonaceous the subsurface (Caldwell et al., 2004). Highly elongated ellipses material has been fully converted to graphite. suggest a degree of structural anisotropy in which horizontal All three conductive regions overlap with mapped carbona- ceous graywackes that have experienced greenschist to amphibo- lite grade metamorphism and metamorphic temperatures from 600 °C to 650 °C(Attoh and Klasner, 1989). Two of the three regions (FA and OA) further overlap AEM surveys with elongate, sub-vertical conductivity anomalies. The FA additionally encloses boreholes that encounter graphite stringers in drill core. I thus con- clude that metamorphic graphite is the source of the intense con- ductivity in the previously recognized Flambeau anomaly, as well as the newly defined Ottawa and Niagara anomalies.

6.2.2. Minnesota To the west of the St. Croix horst, Paleoproterozoic rocks in MN consist of older rocks of the Mille Lacs and North Range Groups exposed in a series of structural panels in east-central MN as well as a broader expanse of younger Animikie Group rocks extending northward into southern Ontario along the north shore of Lake Superior (Fig. 7). The Keweenawan-age Duluth Complex intruded these Paleoproterozoic rocks, dividing Animikie Group rocks along the Lake Superior shoreline into southern and northern exposures in MN and Ontario, respectively. Within the Animikie Group, the Fig. 10. Measured phase tensor ellipses from 20 to 70 s periods at stations located above the conductive Ottawa and Animikie basin anomalies. Ellipses are displayed Virginia and Rove Formations are the most volumetrically signifi- in geographic coordinates with north up. Colored fill indicates the azimuth of the cant, consisting of thousands of meters of graywacke and shale ellipse axes. (For interpretation of the references to color in this figure legend, the (Ojakangas et al., 2001). These formations are considered to be reader is referred to the web version of this article.) 356 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 electrical current flow is direction-dependent, whereas more the high conductivity beneath the MGD to be due to graphite rounded or circular ellipses suggest little or no electrical aniso- formed during metamorphism of Animikie Group rocks that were tropy. Phase tensors also contain directional information, with previously thrust beneath the Mille Lacs and North Range Groups the ellipse axes defining a geo-electric strike direction, which com- near the end of the Penokean orogeny. If Animikie Group rocks monly coincides with the geologic strike direction. Within this con- do not project at depth beneath the McGrath structural panel, it text, the phase tensor ellipses for station above the ABA and the OA is possible that carbonaceous material within older Paleoprotero- are distinct from one another. Those corresponding to the ABA are zoic rocks (Mille Lacs and North Range Groups) is the original rounded and show no consistent direction over the period (depth) source of the inferred graphite. range shown. This is consistent with conductive material deposited in an undeformed basin. The phase tensors corresponding to the 6.3. Lithospheric sutures and Cambrian rifts OA, in contrast, are highly elongated and show a consistent orien- tation of N60°E, sub-parallel to both the regional geologic strike The Archean Great Lakes tectonic zone is discontinuously and the strike of linear conductive anomalies identified in multiple identified geologically and geophysically along its length. It is AEM surveys that overlap the edges of the OA (Fig. 8). This is con- most prominently expressed in magnetic-field data in central sistent with the interpretation of vertical stringers of metamorphic MN (Fig. 2b), potentially extending beneath the Paleoproterozoic graphite aligned with the regional geologic strike and contained rocks along the Malmo Discontinuity (MD, Fig. 9). East of the within highly deformed, steeply dipping structures. MRS in Upper Michigan, the location of the GLTZ can be For all of the reasons above, I conclude that the elevated con- roughly bracketed between exposed rocks of the Superior craton ductivity of the ABA is not due to metamorphic graphite. What is and Minnesota River Valley terrane. The only place where the the cause of the high conductivity within the ABA? The most likely GLTZ is exposed is in the central part of Upper Michigan explanation is metallic sulfides, predominately pyrite and pyrrho- (Sims, 1980). tite. Similar to graphite, sulfides can dramatically reduce the bulk The resistivity model does not directly image the GLTZ, in part resistivity of a rock even if disseminated within a host rock at con- due to overprinting Paleoproterozoic (Penokean) and Mesoprotero- centrations below 1%. Pyrite-containing ores range from 10 O mto zoic (MRS) structures, which have to some extent been localized by 10-4 O m, while pyrrhotite ores range from 10-3 O mto10-5 O m the GLTZ. To first order, the conductive Paleoproterozoic rocks in (Parasnis, 1956). Sulfides are commonly present in reduced car- WI and MI are located south of the GLTZ. In MN, no lateral changes bonaceous marine sediments, the origin of much of the Virginia in resistivity are observed across the magnetically-defined GLTZ, and Rove Formations. A bedded pyrrhotite layer is recognized although a diffuse conductivity anomaly parallels the GLTZ about within the argillite and graywacke of the Virginia Formation, 100 km to the south at 6–9 km depth (Fig. 5b). where it is tens of meters thick in boreholes adjacent to the Duluth In contrast to the GLTZ, the Spirit Lake tectonic zone (NICE Complex (Hauck et al., 1997). While the lateral extent of the bed- working group et al., 2007) is a prominent conductive feature in ded pyrrhotite layer is unknown, this layer, together with the lower-crust (Yang et al., 2015) starting at the western edge sulfide-enhanced rocks throughout the Virginia Formation, are of the station array near the NE/IA border and trending east- taken to be the most likely cause of the high conductivity that is northeast before it is obscured by the MRS near the intersection characteristic of the ABA. of the St Croix and Iowa horsts (Fig. 5b and c). East of the rift, no Wundermann and Young (1987) analyzed a small number of obvious change in conductivity structure is observed. While this wideband MT stations in MN and argue that conductive Animikie boundary has been inferred through central WI from geochemical Group sedimentary rocks extend southward beneath older North data (Van Schmus et al., 2007), at only one isolated location in WI Range and Mille Lacs Groups arranged in fold and thrust belts is there enhanced conductivity comparable (similar conductivity along a south-dipping decollement. The 3D resistivity model is and depth extent) to what is observed along the SLTZ in Iowa. consistent with this interpretation, as the high conductivity of Yang et al. (2015) argue that the SLTZ dips to the north within the ABA is observed to penetrate 30–40 km to the south beneath the lithospheric mantle; the resistivity model presented here does the fold and thrust belt (Fig. 9a), terminating near the north edge not resolve a clearly dipping structure within the mantle, however of the McGrath structural panel. This interpretation of Animikie the crustal conductivity anomaly, most prominent from 10 to Group rocks underthrust beneath panels of Mille Lacs and North 30 km depth, dips to the south-southeast (Fig. 6b), and is consis- Range Group rocks is at odds, however, with two pieces of evi- tently located to the south-southeast of the mapped trace of the dence. First, the Animikie Group is not mapped in fault contact suture as inferred from magnetic-field data (NICE working group, with the thrust belts to the south. Second, an interpretation of 2007). I note that the isolated region of high conductivity in central magnetic data over the area suggests that the thrust belts extend WI (44°N, 90.5°W, Fig. 5c) is located north of the inferred trace of at depth to the north beneath the Animikie basin (Chandler, the SLTZ. Assuming the conductivity anomaly in central WI is 1993). This apparent contradiction can be resolved when the tem- related to the SLTZ, this spatial pattern suggests it may be located poral relationship between late-stage Penokean compression and further north than inferred by Van Schmus et al. (2007). deposition of the Animikie Group is considered. Ojakangas et al. The variable signature of the SLTZ, as imaged in the resistivity (2001) provide evidence that the Animikie Group was laid down model, leads to an alternate interpretation in which the high in a northward migrating foreland basin in front of the developing conductivity sub-parallel to the SLTZ in IA is actually a slice of fold and thrust belt. This suggests that deposition in the Animikie Paleoproterozoic (Penokean) rocks preserved alongside the basin and development of the fold and thrust belt along its south- younger SLTZ. Arguments in favor of this interpretation include ern margin may have been coeval, and that the thrust panels may the unusually high conductivity (throughout the study area such project as a wedge into the Animikie basin, overriding Animikie high conductivity is only associated with Paleoproterozoic Group rocks along a decollement, yet ultimately buried in part metasedimentary rocks), the contradiction between crustal and beneath the youngest Animikie Group sedimentary rocks. lithospheric dips, and a pluton dated at 1810 Ma immediately The conductive anomaly beneath the McGrath gneiss dome north of the SLTZ (Van Schmus et al., 2007) that may be a late- (MGD) differs in that it lies within a structural panel that has stage intrusive associated with the Penokean orogeny. undergone amphibolite grade metamorphism, most likely as part Further south, the transition from Yavapai (1.8–1.7 Ga) to of the mid-Geon 17 event that exhumed the MI/WI gneiss corridor Mazatzal (1.7–1.65 Ga) age crust is defined within the study area (Southwick and Morey, 1991). Given this association, I interpret by a belt of 1.47–1.43 Ga anorogenic plutons (the Green Island P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 357

Plutonic Belt) that stretch through southeast IA, southeast WI, and enhanced conductivity reflects lower-crustal modification that northwest IL (NICE working group, 2007) and is prominent in occurred during the Cambrian breakup of Rodinia. The exception magnetic-field data (Fig. 2b). The extension of this boundary fur- to this may be the Reelfoot Rift, where high conductivity appears ther east and west, as drawn in Figs. 1, 2 and 5, is taken from limited to the thick Phanerozoic section. Whitmeyer and Karlstrom (2007). The conductivity model does not indicate any distinct resistivity signature along this boundary. 6.4. Lithospheric mantle structure Continuing to the southeast, the Nd isotopic boundary between the Mazatzal province to the northwest and younger juvenile crust In examining the mantle beneath the region, I exclude inter- to the southeast (Bickford et al., 2015) is also not evident in the preting regions near or beyond the edge of the station array. This resistivity model. Any distinct resistivity signature between these is done to avoid possible inversion artifacts arising from conduc- provinces has likely been erased by the granite-rhyolite magma- tive regions beyond the station array that the measured data are tism which overprints this boundary. Examining the extent of sensitive to, yet cannot fully constrain with regard to position granite-rhyolite magmatism in relation to the resistivity model, and conductivity. Such problems are more acute at greater model there is some correlation with a region of enhanced conductivity depths, as the long-period MT data used to develop the model at 6–15 km depth (Fig. 5b and c). The origin of this high conductiv- are sensitive to not only deep structure, but also to structure hun- ity is unknown. dreds of kilometers away from any given measurement location. The southeast corner of the station array crosses the Reelfoot The seismically-defined lithosphere-asthenosphere boundary rift–Rough Creek graben–Rome trough intracratonic rift zone (LAB) varies smoothly throughout much of the region and is close (Stark, 1997). These grabens, filled with up to 8 km of Phanerozoic to 200 km depth (Yuan et al., 2014). The depth to the electrically sedimentary rocks, formed during the Cambrian breakup of defined LAB, a transition to higher conductivity thought to reflect Rodinia and the opening of the Iapetus Ocean (Csontos et al., a more hydrous asthenosphere (e.g. Hirth and Kohlstedt, 1996), 2008; Thomas, 1991). The resistivity model reveals linear zones is discussed in Yang et al. (2015) and found to be in close corre- of elevated crustal conductivity along these rift zones, with the spondence with seismic LAB estimates. I focus on the lithospheric highest conductivity at 6–9 km depth along the Reelfoot Rift mantle, and examine a depth slice at 140 km (Fig. 11a). The mantle (Fig. 5b) and from 9–15 km beneath the Rough Creek graben lithosphere is generally resistive (100–1000 X m) beneath the (RCG) and Rome trough (RT, Fig. 5b and c). This east-west trending study area, as expected for stable (depleted and dehydrated) litho- zone of elevated crustal conductivity appears to be continuous sphere. Exceptions to this are four areas outlined in Fig. 11a (white between the RCG and the RT and can be traced for 500+ km at circles). All but one of these areas fall near the edge of the station 37°150N latitude, extending to the isotopic Nd line and the western array and should be treated with a degree of caution. These regions edge of the Granite-Rhyolite Province in southeast Missouri. Taken are an order of magnitude more conductive than the surrounding together, enhanced conductivity traces out a nearly 1000-km long areas with conductivity peaking in the lowermost lithosphere from arcuate path from West Virginia to Missouri, with a southwest 150 to 200 km depth. As this zone is beneath the graphite-diamond trending splay along the Reelfoot Rift. I speculate that the stability field (140 km), I exclude graphite as the cause of the high

Fig. 11. Resistivity model at 140 km depth (a). Underlying hillshaded image is the total magnetic field anomaly. White dashed lines are interpreted terrane boundaries. Black dashed line indicates the trend of MRS. White arrows indicate the approximate SKS splitting direction throughout the region (Frederiksen et al., 2013b). Red circles indicate locations of xenoliths analyzed by Griffin et al. (2004) and discussed within the text. White circles are regions of anomalous conductivity discussed in the text. A sampling of resistivity-depth functions for regions A (red) and B (black) are compared to the unified olivine resistivity model of Gardés et al. (2014) (thick black lines) with varying concentrations of water in olivine (b). The lithosphere-asthenosphere boundary (LAB) is shown as a thick gray line. Thick green line is the lithospheric geotherm (Hasterok and Chapman, 2011) used to convert depth to temperature. Below 250 km the model is affected more by regularization than by measured data. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) 358 P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 conductivity, although it may contribute at shallower depths surements on hydrous olivine. I couple this conductivity model to a within the lithosphere. conductive geotherm for the region (green line, Fig. 11b), account- The spatial pattern of resistivity in the deep lithosphere bears ing for variable heat production and surface heat flow of 40 mW/ little resemblance to the configuration of the MRS system in the m2 (Hasterok and Chapman, 2011). crust (dashed black line in Fig. 11a) or the Archean and Proterozoic Fig. 11b shows this model for a dry lithospheric mantle and for terrane boundaries (dashed white lines) that trend southwest- mantle with water concentrations of 40 and 300 wt ppm in olivine. northeast throughout the region. Likewise, there appears to be lit- Throughout much of the study area, the resistive deep lithosphere tle association between regions of conductive lithosphere and the is consistent with olivine water contents between 0 and 40 wt extent of granite-rhyolite magmatism (compare to Fig. 1). Finally, ppm, in accordance with thick stable lithosphere that has not expe- the mantle resistivity structure does not mimic seismic SKS split- rienced recent tectonism. In contrast, the lithosphere within the ting directions (range denoted by white arrows), which reflect conductive regions (A) require olivine water contents of 300 wt lithospheric anisotropy and are subparallel to the direction of ppm throughout the lower half of the lithosphere. absolute plate motion (Frederiksen et al., 2013b). Based upon the current understanding, the required water con- What do the resistivity variations in the mantle lithosphere rep- tent in these regions is incompatible with the presence of a stable resent? The conductivity of continental lithosphere is to first order lithosphere. This apparent contradiction has been observed by controlled by temperature. However, the entire survey area, save others (e.g. Fullea et al., 2011; Selway et al., 2014) and may stem for the southeast corner, has not undergone significant tectonic from discrepancies between various laboratories making olivine modification since Grenville times. Hence, it is unlikely that spatial conductivity measurements or from water in other mantle miner- variations in lithospheric geotherm can account for the variability als not taken into account by the current models. While acknowl- in mantle resistivity. Compositional variations are also unlikely to edging this issue of scale (too much water to maintain a stable produce the observed variations. Jones et al. (2009) and Selway lithosphere, too high a model conductivity to be explained by mod- et al. (2014) both demonstrate the minor impact of composition est water content), I consider hydration to remain the most plausi- on lithospheric mantle conductivity, finding, for example, the con- ble explanation for the enhanced conductivity of the SCLM. ductivity difference between harzburgite and iron-rich lherzolite Assuming that water within nominally anhydrous mantle minerals to be at the level of MT measurement uncertainty. (olivine, orthopyroxene, clinopyroxene) is the driver behind the In the absence of an unknown conductivity mechanism, the variations in lithospheric resistivity, is there evidence for metaso- most likely explanation for the observed conductivity variations matized mantle beneath the mid-continent? is dissolved water in nominally anhydrous mantle minerals. Cryptic metasomatism is the result of rising or percolating Olivine, the major component in mantle peridotite, can accommo- melts interacting with surrounding mantle peridotite. Due to their date water at concentrations up to a few hundred parts per million inherent ability to permeate a peridotite matrix, some of the most (ppm) (Bell and Rossman, 1992). Laboratory data show significant effective metasomatizing agents are volatile-rich silicate or car- enhancement in conductivity for olivine, orthopyroxene, and gar- bonatite melts (Coltorti and Grégoire, 2008; Luth, 2003). As the net at low temperatures and modest water content (Poe et al., putative agent behind the last tectonic event to affect the region, 2010; Dai and Karato, 2009; Yoshino et al., 2009). These hydrated the Keweenawan mantle plume is considered the most likely minerals can have conductivities that are significantly greater than source of mantle metasomatism. However, even small degrees of their dry mineral counterparts, although the degree of conductivity partial melting will deplete the lithospheric mantle. Thus, explain- enhancement, particularly at asthenospheric temperatures, is the ing the modeled conductivity variations in terms of metasomatism subject of some debate (Gardés et al. (2014) and references requires that either lithospheric melting due to the plume was pat- therein) and decreases considerably near the base of the litho- chy and did not affect the entire lithosphere, or that the final stage sphere (Evans, 2012). of rift magmatism was from magmas of plume origin (i.e., no litho- Beneath the stable mid-continent, the thick lithospheric mantle spheric mantle component). has persisted for over a billion years. The stability of sub- Nicholson et al. (1997) provide support for the latter, docu- continental lithospheric mantle (SCLM) is in fact dependent upon menting a rift-wide, time-progressive change in the source of rift its relative state of dehydration, making it viscous relative to the basalts. While earlier basalts indicate significant interaction with underlying asthenosphere, and resistant to delamination. mantle lithosphere, magmas from a latter voluminous phase are Extensive melting during formation and stabilization is considered dominantly of plume origin, and the youngest magmatic products to have depleted the SCLM of most of its water. Given that show no lithospheric component, reflecting a mixture of melts increased water content in olivine is accompanied by a decrease from the plume and a depleted asthenospheric mantle source. in viscosity, it is reasonable to expect the olivine water content The absence of lithospheric melting during this final magmatic of SCLM to be less than that of the asthenosphere, estimated to phase suggests that melt-metasomatism last affected the dissolved be around 100 ppm (Dai and Karato, 2009). Studies of SCLM xeno- water content of the lithosphere. liths by Peslier et al. (2010) suggest the water content of SCLM is Independent support for metasomatized mantle in two of the typically <100 ppm, although it may vary throughout the litho- four conductive regions outlined in Fig. 11a comes from Griffin spheric column. Recent MT studies of the SCLM, however, in com- et al. (2004), who analyzed the compositional makeup of the man- bination with laboratory conductivity models, require significantly tle lithosphere throughout North America based upon major and higher olivine water contents to fit estimates of lithospheric con- trace element analyses of garnet in mantle xenocrysts. Within ductivity (e.g. Fullea et al., 2011; Selway et al., 2014). the model area, xenoliths from WI and Upper MI (red circles in Inverted resistivity-depth models from both the conductive (A) Fig. 11a) are dominated by signatures of fertile or melt- and resistive (B) regions (Fig. 11b) show a general decrease in metasomatized mantle from depths of 100–175 km. The signature resistivity with increasing depth from the Moho to the LAB. Within of melt-metasomatism is strongest in the WI samples and the conductive regions, however, the lithospheric mantle is increases with depth, especially below 175 km. Mantle xenocrysts roughly an order of magnitude more conductive than in adjacent in kimberlites from KS are also dominated by a signature of melt- regions. I compare these inverted resistivity-depth models to the metasomatism, an observation that Griffin et al. (2004) attribute to unified hydrous conductivity model of Gardés et al. (2014),a asthenospheric upwelling related to the MRS. model for olivine conductivity as a function of temperature and I conclude that variations in lithospheric mantle conductivity water content calibrated against a diverse suite of laboratory mea- beneath the study area reflect variations in the degree of mantle P.A. Bedrosian / Precambrian Research 278 (2016) 337–361 359 hydration. Plume-dominated magmas comprising the final phase Allen, D.J., Hinze, W.J., 1992. Wisconsin gravity minimum: solution of a geologic and of MRS magmatism are considered to be the most-likely metasom- geophysical puzzle and implications for cratonic evolution. Geology 20, 515– 518. http://dx.doi.org/10.1130/0091-7613(1992) 020<0515:WGMSOA>2.3. atizing agent. Why some areas remain depleted (resistive) while CO;2. others are fertile (conductive) is an interesting avenue for further Allen, D.J., Hinze, W.J., Dickas, A.B., Mudrey, M.G., 1997. Integrated geophysical research. I note in conclusion that the melt-metasomatized region modeling of the North American Midcontinent rift system: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota. beneath WI coincides roughly with the extent of both the Wiscon- Geol. Soc. Am. Special Pap. 312, 47–72. sin magmatic arc terranes and the voluminous 1.47 Ga Wolf River Anderson, J.L., 1983. Proterozoic anorogenic granite plutonism of North America. In: batholith, the latter interpreted by Allen and Hinze (1992) to have Geological Society of America Memoirs. Geological Society of America, pp. 133– 154. played a role in focusing the Mid-continent rift system. Anderson, R.R., 2006. Geology of the Precambrian Surface of Iowa and Surrounding Area (Iowa Geological Survey Open File Map No. OFM-06-7). 7. Conclusions Anderson, R.R., 1997. Keweenawan Supergroup Clastic Rocks in the Midcontinent Rift of Iowa. Geological Society of America Special Papers 312, 211–230. Anderson, R.R., 1992. The Midcontinent Rift of Iowa (Ph.D.). University of Iowa. Over a protracted period in the Proterozoic, the North American Anderson, R.R. (Ed.), 1990. The Amoco M.G. Eischeid #1 Deep Petroleum Test, mid-continent was assembled in much the same way that the west- Carroll County, Iowa: Preliminary Investigations. Iowa Department of Natural ern continental margin developed in the Mesozoic and Cenozoic. Resources Geological Survey Bureau, Special Report Series no 2. Anderson, R.R., McKay, R.M., 1989. Clastic Rocks Associated with the Midcontinent Unique to the mid-continent is the preservation of this tectonic col- Rift System in Iowa (U.S. Geological Survey Bulletin No. 1989-I). lage, largely unmodified by tectonic events since the 1.1 Ga Mid- Attoh, K., Klasner, J.S., 1989. Tectonic implications of metamorphism and gravity continent rift system nearly tore it apart. I have investigated aspects field in the Penokean Orogen of northern Michigan. Tectonics 8, 911–933. http://dx.doi.org/10.1029/TC008i004p00911. of the structure and evolution of the region through the lens of a 3D Behrendt, J.C., Green, A.G., Cannon, W.F., Hutchinson, D.R., Lee, M.W., Milkereit, B., resistivity model based upon EarthScope magnetotelluric data. The Agena, W.F., Spencer, C., 1988. Crustal structure of the Midcontinent rift system: geometry, extent, and segmentation of the MRS was examined in results from GLIMPCE deep seismic reflection profiles. Geology 16, 81. http://dx. doi.org/10.1130/0091-7613(1988) 016<0081:CSOTMR>2.3.CO;2. detail and shed light on the structural complexity and multi-phase Behrendt, J., Hutchinson, D., Lee, M., Thornber, C., Trehu, A., Cannon, W., Green, A., development of the Lake Superior graben. An unrecognized expanse 1990. GLIMPCE seismic reflection evidence of deep—crustal and upper-mantle of (concealed) Precambrian deltaic deposits in Kansas was identified intrusions and magmatic underplating. Tectonophysics 173, 595–615. and speculated to result from axial drainage along the southwest rift Bell, D.R., Rossman, G.R., 1992. Water in Earth’s mantle: the role of nominally anhydrous minerals. Science 255, 1391. arm akin to the modern Rio Grande delta. A prominent conductor Berendsen, P., 1997. Tectonic evolution of the midcontinent rift system in Kansas. attributed to Cambrian rifting in Arkansas, Missouri, Tennessee Geol. Soc. Am. Special Pap. 312, 235–242. and Kentucky suggests rifting was more extensive than previously Beyssac, O., Goffé, B., Chopin, C., Rouzaud, J.N., 2002. Raman spectra of carbonaceous material in metasediments: a new geothermometer. J. Metamorph. Geol. 20, thought. The highest conductivity within the mid-continent is 859–871. imaged in Minnesota, Michigan, and Wisconsin where it is coinci- Bickford, M.E., Van Schmus, W.R., Karlstrom, K.E., Mueller, P.A., Kamenov, G.D., dent with Paleoproterozoic metasedimentary rocks. The high con- 2015. Mesoproterozoic-trans-Laurentian magmatism: a synthesis of continent- wide age distributions, new SIMS U-Pb ages, zircon saturation temperatures, ductivity is attributed to metallic sulfides, and in some cases, and Hf and Nd isotopic compositions. Precambr. Res. 265, 286–312. http://dx. graphite. 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Closing of the Midcontinent rift—a far-field effect of Grenvillian provided by Ray Anderson, Ryan Clark, Bill Cannon, Dick Ojakangas compression. Geology 22, 155–158. and Laurel Woodruff. Thanks to Ben Bloss and Emily Hart and for Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., Spencer, C., 1989. performing lab conductivity measurements. This work would not The North American Midcontinent Rift beneath Lake Superior from GLIMPCE have been possible without the efforts of the National Geoelectro- seismic reflection profiling. Tectonics 8, 305–332. http://dx.doi.org/10.1029/ magnetic Facility at Oregon State University, Zonge Engineering TC008i002p00305. Cannon, W.F., Hinze, W.J., 1992. Speculations on the origin of the North American and Research Organization, and Green Geophysics, all of whom Midcontinent rift. Tectonophysics 213, 49–55. http://dx.doi.org/10.1016/0040- oversaw or carried out the acquisition of the EarthScope MT data 1951(92)90251-Z. presented here. Adam Schultz, Gary Egbert, and Anna Kelbert are Cannon, W.F., Lee, M.W., Hinze, W.J., Schulz, K.J., Green, A.G., 1991. Deep crustal particularly thanked for the data QA/QC and time-series processing structure of the Precambrian basement beneath northern Lake Michigan, midcontinent North America. Geology 19, 207–210. which make this data set accessible to the earth science community. Cannon, W.F., Ottke, D., 1999. Preliminary digital geologic map of the Penokean Constructive reviews by Bill Cannon, Ray Anderson, Rob Evans, and (Early Proterozoic) continental margin in northern Michigan and Wisconsin (U. Randy Keller greatly improved this manuscript. S. Geological Survey Open-File Report No. 99-547). Cannon, W.F., Peterman, Z.E., Sims, P.K., 1993. 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