Abstract

We present structural data and cross sections from four transects that together cover much of the Eastern Belt of the Franciscan accretionary complex. The westernmost, Middle Eel transect includes jadeite-lawsonite facies rocks (the Taliaferro Metamorphic Complex, TMC) intercalated with lawsonite-albite facies metagreywacke. The TMC shows subduction-related imbricate thrusting, refolded by upright folds with amplitudes of 100–1500 m, and it is cut by abundant normal faults that contributed to exhumation. East of the Coast Ranges divide, three linked transects along Thomes Creek cover the transition from lawsonite-albite facies metagreywackes to the blueschist facies South Fork Mountain Schist. The western section shows thick-bedded metagreywacke intercalated with broken formation, deformed by NW-vergent folds and associated thrusts and pressure-solution cleavage, and intensively dissected by abundant low-angle normal faults. In the central section, thin-bedded metagreywacke, broken formation, and conglomerate show an early foliation and folds overprinted by E-vergent folds and crenulation cleavage. The South Fork Mountain Schist forms the easternmost section and records the most intense deformation. The dominant foliation is a differentiated crenulation cleavage that has been refolded by NW vergent folds with amplitudes of millimeters to hundreds of meters. Structural relationships in the South Fork Mountain Schist exposed in Cottonwood Creek farther north are similar to those in Thomes Creek, indicating that our observations have regional significance. All the contractional structures and ductile deformational fabrics in these transects formed under high-P low-T metamorphic conditions during subduction and accretion, and the dominant deformation mechanism was pressure solution. Exhumation was achieved primarily by intensive normal faulting on the outcrop scale, and normal sense motion on the Coast Range fault. This paper provides the first documentation of syn-subduction normal faulting within the Franciscan Complex.

INTRODUCTION

The internal structure of accretionary complexes is poorly known and understood: most active examples are largely or completely submerged, and ancient examples are commonly strongly modified by later events such as continent or arc collision. The problem is compounded by the fact that accretionary complexes are commonly largely composed of relatively monotonous greywacke sandstone and shale sequences, without well-developed lithostratigraphy, and with complicated and disruptive structural styles. This hinders field investigations of emergent examples as well as seismic studies of currently active complexes in submerged fore-arcs. In spite of this, some excellent seismic studies have demonstrated that the frontal regions of accretionary wedges are dominated by imbricate thrusting (Davey et al., 1986; Davis and Hyndman, 1989; Moore et al., 1990; Morgan and Karig, 1995), and this has been confirmed by detailed studies of some well-exposed emergent examples (Moore and Karig, 1980; Wahrhaftig, 1984; Platt et al., 1988; Meneghini and Moore, 2007, Wakabayashi, 2017). The structure of the more deeply buried interiors of accretionary wedges is less well documented, and in particular the processes driving exhumation of these rocks remain controversial. The structure in the base of the accretionary wedge is likely to be dominated by the process of subcretion or underplating, which may produce large-scale thrust-sheets (Wakabayashi, 1992) or duplexes (Kimura et al., 1996). Exhumation has variously been attributed to return flow in a subduction channel (Cloos and Shreve, 1988), wedge extrusion (Maruyama et al., 1996), normal faulting in the upper part of the accretionary wedge, or in the overlying fore-arc basin (Platt, 1986; Jayko et al., 1987; Harms et al., 1992; Wakabayashi and Unruh, 1995; Constenius et al., 2000; Schemmann et al., 2008; Unruh et al., 2007), or erosion (Feehan and Brandon, 1999; Ring and Brandon, 1999; Ring, 2008).

The internal structure of the Franciscan Complex in California is particularly poorly known, in part because of the abundance of highly disrupted rocks generally referred to as mélange, and in part because of poor exposure. The aim of this paper is to present detailed field relationships along a transect across the relatively coherent eastern belt of the Franciscan in well-exposed river sections in the northern Coast Ranges, and to discuss the significance of the structure in terms of subduction, underplating, and exhumation.

GEOLOGIC SETTING

The Franciscan Complex is the archetypal accretionary complex formed at a convergent plate boundary (Bailey et al., 1964; Ernst, 1970; Wakabayashi, 1999), and reflects the subduction of tens of thousands of km of oceanic lithosphere along the western margin of North America from mid-Jurassic to mid-Tertiary time. The oceanic lithosphere carried with it seamounts and a pelagic sedimentary cover, and at times accumulated great thicknesses of clastic sediment in a trench environment, and much of this material was scraped off and accreted to form the accretionary wedge. Some was accreted at shallow depths near the trench, but some was carried beneath the wedge, and underplated at depths of 20–40 km, where it was metamorphosed under high-pressure–low-temperature (high-P/low-T) conditions (Ernst, 1971). Overall, the Franciscan comprises ∼80% greywacke sandstone and shale, the remainder being predominantly mafic volcanic rocks and minor amounts of radiolarian chert and pelagic limestone (Bailey et al., 1964). Perhaps as much as 30% of the outcrop area of the Franciscan shows a block-in-matrix texture, generally referred to as broken formation or mélange (Hsü, 1968). The matrix is commonly shale, usually with a scaly fabric; the blocks, which may vary in size from a few mm to tens or even hundreds of m, consist mainly of greywacke sandstone in broken formation, but may include volcanic rocks, chert, eclogite, garnet amphibolite, and blueschist, in which case the rock is referred to as mélange. Mélange and broken formation have variously been interpreted as olistrostromes, mass flows, or debris flows of sedimentary origin (Cowan, 1985; Wakabayashi, 2011, 2015; Platt, 2015), or as a result of tectonic processes such as return flow in the subduction channel (Cloos, 1982).

The Franciscan Complex is bounded to the east by the mid-Jurassic Coast Range ophiolite. This represents the oceanic crust on the deformed leading edge of the North American plate (Hopson et al., 2008), and is overlain by fore-arc basin sediments of the Great Valley Group (Dickinson et al., 1996). Neither the Coast Range ophiolite nor the Great Valley Group shows significant metamorphism. They are separated from the Franciscan Complex by the Coast Range fault, which dips steeply E along the eastern margin of the Franciscan, but scattered outliers of Coast Range ophiolite and Great Valley Group across the northern Coast Ranges suggest that it is regionally gently dipping (see Wakabayashi, 2015, for a review). The Coast Range fault was originally identified by Ernst (1970) as the Mesozoic paleo-subduction zone, but it is now generally recognized that the fault has been cut or reactivated by later normal sense motion, allowing exhumation of the underlying subduction complex (Platt, 1986; Jayko et al., 1987). We note that Ring and Brandon (1994, 1999) and Ring (2008) have argued that exhumation of the Franciscan Complex was accomplished by erosion of an emergent fore-arc high, and that the Coast Range fault is a later out-of-sequence thrust. The consistency of hanging wall and footwall rock sequences along the Coast Range fault suggest, however, that it still closely approximates the original subduction zone contact.

The Franciscan in the northern California Coast Ranges has traditionally been divided into three belts, based on the timing of accretion, grade of metamorphism, and the overall structural style (Fig. 1) (Berkland et al., 1972; Ernst, 1975). The character and boundaries of these belts are loosely defined and somewhat controversial, but they provide a useful framework for discussion of the internal structure of the Franciscan. The Eastern Belt is the oldest, shows widespread high-P/low-T (lawsonite-albite and blueschist-facies) metamorphism, and a relatively coherent structural style. The Central Belt shows a very disruptive structural style, containing large volumes of scaly-clay mélange with a low metamorphic grade (lawsonite-albite and prehnite-pumpellyite facies), but with relatively abundant tectonic blocks of eclogite, garnet amphibolite, and blueschist. The Coastal Belt is the youngest, includes significant tracts of deformed but coherent siliciclastic sedimentary rock, and is essentially unmetamorphosed (zeolite facies).

The Eastern Belt

The Eastern Belt of the Franciscan includes large tracts of lawsonite-albite facies siliciclastic rocks with a fairly coherent structure, some substantial bodies of mélange, and a number of large sheets or slabs of blueschist-facies metasediment and metabasalt (Suppe, 1973; Brown and Ghent, 1983; Bröcker and Day, 1995). The largest of these thrust sheets is the South Fork Mountain Schist (Blake et al., 1967), which is several km thick and extends ∼250 km along strike on the eastern margin of the Franciscan. This is the largest coherent body of blueschist-facies rock in the northern Coast Ranges, and is characterized by a very strong schistose fabric in both metapelitic and metabasaltic rocks (Blake et al., 1967).

Suppe (1973) mapped a large area of the eastern belt, encompassing most of the area discussed in this paper. He distinguished two broad “facies” within the siliciclastic rocks: predominantly thick-bedded greywacke sandstones and shales, and scaly clay mélange or broken formation. He also identified a fault-bounded sheet of jadeite-bearing blueschist-facies metasediments and metabasalt, which he named the Taliaferro metamorphic complex. This lies to the west of, and appears to be distinct from, the South Fork Mountain Schist, and appears to be intercalated with lower-pressure lawsonite-albite facies rocks. Suppe identified the lower tectonic boundary of the South Fork Mountain Schist as a major thrust, the Log Spring thrust, separating it from lower grade rocks beneath it. Suppe’s cross sections show the South Fork Mountain Schist and its basal thrust dipping steeply east beneath the Coast Range fault, but flattening westward, so that they intersect the topography near the crest of the Coast Range (Fig. 2).

Subsequent mapping (e.g., Worrall, 1981; Blake and Jayko, 1983) led to the distinction of a number of lithotectonic units, made up of greywacke sequences, mélange, or broken formation, separated by major faults, and Jayko and Blake (1989) then subdivided the Eastern Belt as a whole into the Pickett Peak terrane (which includes the South Fork Mountain Schist) and Yolla Bolly terranes. These were distinguished in part by their primary lithological characteristics, and partly by their textural grade, which refers to the intensity of fabric development and the degree of recrystallization of clastic sedimentary grains in the greywackes. There is clearly an overall increase in metamorphic and textural grade eastward and structurally upward across the Eastern Belt (see also Blake et al., 1967; Suppe, 1973), but it is not clear whether this is an appropriate basis for dividing the belt into separate terranes. As noted below, we were not able to confirm the locations shown by Jayko and Blake (1989) for some of the major faults and unit boundaries, and for this reason we do not use their terrane classification in this paper. Geothermometry on rocks from the Eastern Belt using laser Raman spectrometry on carbonaceous material is in progress, and may help in providing a clearer definition of the tectonic units.

Detrital zircon dating (Dumitru et al., 2010) has established that protolith ages of the clastic sedimentary rocks in the Eastern Belt are Early Cretaceous (137–111 Ma); the South Fork Mountain Schist is likely to be the oldest. Ar-Ar ages on white mica from the South Fork Mountain Schist are around 121 Ma; these may be crystallization or cooling ages, but in either case are likely to be close to the time of accretion (Dumitru et al., 2010). The timing of exhumation is less certain, but Eastern Belt rocks are likely to have cooled through the apatite fission-track annealing window (corresponding to a depth of ∼10 km) by Late Cretaceous time (Dumitru, 1989; Tagami and Dumitru, 1996).

The eastern boundary of the Franciscan in the northern Coast Ranges shows some complexity. The South Fork Mountain Schist lies in tectonic contact with slivers of low-grade volcanic and siliciclastic rocks, which may be correlative with the mid-Jurassic Galice Formation of the Klamath Mountains (Jayko and Blake, 1986). These low-grade rocks are overlain in turn by a highly disrupted assemblage of ophiolitic rocks, which has been referred to as the Tehama-Colusa mélange (Hopson and Pessagno, 2004; Shervais et al., 2011). The Tehama-Colusa mélange lies in the same structural position relative to both the Franciscan and the overlying Great Valley sequence as the Coast Range ophiolite, but its affinities are controversial, and it lies in fault contact with the Great Valley sequence, which is tilted steeply east near the boundary. The lower part of the Great Valley sequence is cut by several large faults, including the Paskenta fault. These are sinistral in their current orientation, but may originally have been normal faults prior to late Cenozoic tilting (Constenius et al., 2000), and appear to sole onto the Coast Range fault.

OVERVIEW OF THIS STUDY

In common with much of the Franciscan Complex, exposure in the northern Coast Ranges is for the most part very poor, and is largely restricted to road-cuts, rivers, and the summit regions of some of the higher mountains. This, combined with the lack of established stratigraphy, and the complexity and intensity of the deformation, explains the uncertainties in the geological relationships, as discussed above. The river sections, however, can provide good to outstanding exposure over long distances, although they are difficult to access. This study has been confined to the river sections, as they provide the only way to carry out detailed analysis and correlation of deformational structures.

We present a detailed cross section and structural analysis of the rocks of the Eastern Belt along roughly E-W–trending segments of two major drainages: Thomes Creek on the eastern side of the Coast Ranges, and the Middle Fork of the Eel River on the western side of the main divide. A shorter section in Cottonwood Creek, to the north of Thomes Creek, was also investigated in less detail, to check whether the structures observed in Thomes Creek are regionally consistent and not local or anomalous. While these sections do not provide a continuous transect, we believe that they provide sufficient information for us to make well-substantiated statements about the character of the protolith assemblage, the origin of the broken formation, and the structures associated with subduction, accretion and exhumation. The locations of these detailed sections are shown on a regional map and cross section (Figs. 1 and 2). After a description of the two sections, we integrate our observations to interpret the deformational history of the Eastern Belt.

THE MIDDLE EEL SECTION

The section (Fig. 3) extends ∼4 km NE along the Middle Fork of the Eel River from the bridge where forest road FH7 from Covelo to Elk Creek crosses the river. The average orientation of the section is approximately normal to the strike of the structures, and the data have been projected onto a section line that is 3 km long. The section includes rocks of the Taliaferro Metamorphic Complex (TMC), together with lawsonite-albite facies greywacke sandstone, shale, and broken formation, which lie structurally below it. We also constructed a short section through the TMC farther north, in Beaver Creek, which is a small tributary to the Eel River. The Beaver Creek section provides useful additional information on the structure and stratigraphy of the TMC. Structural data from the Middle Eel sections is presented in Figure 4.

The Taliaferro Metamorphic Complex in the Middle Eel River forms a tight synform 1.5 km across, with a clearly defined internal stratigraphy and structure. The stratigraphy consists of massive metabasaltic rocks metamorphosed in blueschist facies, overlain by metachert with pelitic interlayers, followed by dark organic-rich phyllite and greywacke sandstone. In Beaver Creek there is a perfectly exposed section from metabasalt through metachert with deformed radiolarians, followed by a rapid transition into metagreywacke (Fig. 5A).

Metabasalt carries the assemblage glaucophane + lawsonite + sphene (Fig. 6A); primary igneous textures have been largely destroyed, and it has a variably developed deformational fabric. It is internally imbricated, and on the NE limb of the major synform in the Middle Eel there are four thrust imbrications, each one repeating the chert-shale section above the basalts (Fig. 5B). The slices dip and young SW on this limb of the synform, and kinematic data from the thrusts indicate NE-directed thrusting (Fig. 4B). On the SW limb of the synform only two imbrications are identifiable: these young and dip to the NE, but we found no kinematic data. The orientations of the faults and the slip lineations are likely to have been modified by the folding, and given the NE direction of motion identified on the NE limb, it is likely that the slip direction on the SW limb is down-dip in its present orientation. The synform itself trends NW-SE, but must be a younger structure. Stretching lineations in the metachert layers trend around E-W (Fig. 4A).

The graphitic phyllite that lies above the basalt-chert section carries chlorite + white mica + lawsonite + jadeitic pyroxene. It has a strong stylolitic pressure-solution cleavage, and carries numerous sheeted quartz veins that are tightly folded and transposed parallel to the cleavage (Fig. 5C). These veins are particularly well developed along and close to the thrust faults that repeat the metabasaltic section (Fig. 5D). The veins may form up to 50% by volume of the rock, and reflect the precipitation of quartz dissolved during pressure solution, testifying to the importance of dissolution-precipitation creep during subduction and accretion. Some of these veins contain coarse prismatic jadeite and lawsonite (Fig. 6B), demonstrating that they formed at the time of peak pressure. The overlying greywacke sandstones in Beaver Creek have relict detrital textures, and carry lawsonite and fibrous jadeite replacing detrital sodic plagioclase (Fig. 6C). All these rocks are locally deformed by one or more sets of small-scale folds with axial-plane crenulation cleavage.

The lower boundary of the TMC is well exposed on the north limb of the synform, where metabasalt lies in fault contact with lawsonite-albite facies greywacke sandstone and shale beneath. This contact is marked by an extensive sheeted vein complex, and brecciation and veining of the metabasalt. It must be a post-metamorphic thrust, as it places higher pressure TMC rocks above lower pressure lawsonite-albite facies metasediments. On the south limb of the synform the contact has been cut or reactivated by normal faulting (see below). The upper boundary of the TMC is not exposed in the Middle Eel section, but is exposed in Beaver Creek, where it is defined by an array of normal sense shear zones (see below).

The TMC as a whole, together with the lawsonite-albite facies rocks that lie below it, are folded on scales of 100–1500 m by NW-trending, approximately upright folds. These structures may not all have formed at the same time. The major synform in the TMC clearly refolds the imbricate structures within it, and appears to predate the widespread normal faulting (see below). This suggests that it occurred late in the accretionary process, or early during exhumation, as a mechanism of continued shortening and thickening of the underplated rocks.

Both the TMC and the lawsonite-albite facies rocks around it are affected by intensive normal faulting. These faults define many of the present-day boundaries between different rock units, and clearly cut across layering and foliation in many locations. The lower boundary of the metabasalts on the SW side of the major syncline in the TMC is a significant NE-dipping normal fault, which crosscuts or reactivates an earlier thrust that placed the metabasalts above graphitic phyllite (Fig. 3). A set of SW-dipping normal faults offset graphitic phyllite above the metabasalts on the same limb of the syncline (Fig. 7A). These are distinctive because the faults are occupied by quartz veins that carry lawsonite, but not jadeite; the sheeted veins in the graphitic phyllites, however, carry both lawsonite and jadeite (Fig. 6B). This suggests that the normal faulting was accompanied by significant decompression. Normal faults cut and offset the imbricate thrusts on the NE limb of the syncline (Fig. 5B), and an array of conjugate normal-sense shear zones defines the upper boundary of the TMC in Beaver Creek (Figs. 4D, 6D, and 7B). Where normal faults cut shaly or slaty rocks, in both the TMC and the lawsonite-albite facies rocks, they create a scaly-clay fabric with shear bands (Fig. 6D), whereas in the greywackes the faults are discrete and commonly marked by quartz veins. Normal faults mostly have gentle dips, very variable slip directions, and commonly occur as conjugate sets. The majority dip NE, with a mean dip of 28/041, and a mean slip direction of 22/089, with a very large scatter from SE to NE (Fig. 4C). A smaller number of W or NW dipping normal sense shears have a NW sense of slip.

A few minor strike-slip faults are present. Dextral faults strike 300–330; sinistral faults strike 084–116. These are reasonably interpreted as minor structures related to faults of the San Andreas system.

The lawsonite-albite facies rocks lying structurally below the TMC comprise interlayered dark organic rich pelite and massive thick-bedded greywacke sandstone, together with significant bodies of broken formation. The greywacke sandstones generally show low internal strain, well-developed detrital textures, and a weak disjunctive cleavage produced by pressure-solution (Ring and Brandon, 1999; Bolhar and Ring, 2001) (Fig. 7D). The pelites have clearly accommodated most of the strain, and generally show a simple slaty cleavage, at a low or moderate angle to bedding, that appears to be axial planar to the large-scale folds. Much of it, however, shows a scaly-clay fabric indicative of deformation under semi-brittle conditions, associated with the abundant faulting. The broken formation consists largely of fragmented beds of greywacke sandstone floating in a dark pelitic matrix. The sandstone fragments are commonly aligned, defining a fabric that may be folded (Fig. 7E). Locally, there are small, irregular bodies of metavolcanic rock, which commonly show a fragmental texture (Fig. 7F). These may have been emplaced as individual slides or olistostromes of pyroclastic material.

As noted above, the TMC is bounded below by a thrust, which must have been active after peak metamorphic conditions were reached in the TMC, and in Beaver Creek it is bounded above by an array of normal faults. Its present structural position, as a slice bounded above and below by lower-grade rocks, appears to have resulted from two separate tectonic processes: late-stage thrusting during the underplating process, and normal faulting during exhumation.

THOMES CREEK

The Thomes Creek section is conveniently divided into three parts, with different structural characteristics, which are described separately here. The western, upstream part (Slab section) comprises lawsonite-albite facies metasediments lithologically similar to those in the Middle Eel section. The central part (Lanz section) is made up of a similar sequence with a better-developed cleavage, and with a fairly systematic set of late folds associated with a crenulation cleavage. The eastern (downstream) part corresponds to the South Fork Mountain Schist. These three sections are separated by prominent bodies of metabasaltic rocks, which may lie along major thrust boundaries. The Slab section is structurally lowest (Fig. 2), with the Lanz and South Fork Mountain Schist sections lying progressively higher in the sequence. The Slab section is separated from the Middle Eel section described above by a broad zone of poor exposure along the crest of the Coast Ranges. According to the Willows quadrangle map (Blake et al., 1992) this area is occupied by mélange, together with a body of TMC rocks several km across. Given the overall easterly dip of the Franciscan (Fig. 2), these rocks probably lie structurally below the Slab section described here.

Thomes Creek: Slab Section

The Slab section (Fig. 8), named after the concrete slab where Forest Road 24N01 crosses Thomes Creek, consists of thin- to thick-bedded greywacke sandstone and shale sequences and numerous thick beds of pebble conglomerate, all interlayered on various scales with broken formation. Sandstone beds locally show grading, with sharp bottoms and diffuse tops, and the thicker beds commonly contain abundant shale rip-up clasts concentrated near the top of the bed. Sandstone dikes are common (Fig. 9A), indicating fluid overpressuring prior to lithification. Interlayered broken formation is clearly stratigraphically bound, with a crude alignment of sandstone clasts parallel to the bedding in the surrounding rocks (Fig. 9B), and is locally quite strongly folded (Fig. 9C). Clast morphology is generally irregular with rounded edges and re-entrants, suggesting that the sand was unlithified at the time of disruption (Fig. 9C), but locally forms bedding-parallel slabs with planar surfaces that may have been partly lithified.

The sequence is folded on scales from a few cm to several hundred meters (Figs. 8, 10A, and 10B), and bedding is highly variable in orientation. Folds trend NE-SW, are predominantly N- to NW-vergent, and are accompanied by a moderately well-developed slaty cleavage in pelites, and a spaced pressure-solution cleavage in sandstones (Figs. 10C and 10D). Some of the folds are accompanied by minor SE-dipping thrust faults. The cleavage generally dips moderately SE, but has locally been tilted into a NW dip: these variations may be a result of the extensive disruption by the later faults that cut the sequence (see below).

The greywacke sandstones show relict detrital textures and a simple pressure solution cleavage, accompanied by crystallization of quartz and white mica in pressure shadows (Fig. 10C). Pelites carry a simple slaty cleavage. This is clearly overgrown by small lawsonite tablets (Fig. 10D), indicating that it formed early in the subduction history. Both rock types locally show sheeted vein complexes parallel to the cleavage, some of which carry abundant aragonite, partly altered to calcite. The folds and cleavage may therefore reasonably be attributed to early stages in the subduction process, most likely contraction associated with underplating.

At the north end of the section, changes in younging direction revealed by grading in massive NW-dipping greywacke sandstone beds appear to predate the visible folding and the S-dipping cleavage (Fig. 8). This may be a result of early folding in a nonmetamorphic environment at the wedge front, or soft-sediment slump folding.

In the center of the section a zone of broken formation a few tens of meters thick is strongly deformed and foliated, in contrast to the rest of the sequence, which shows only a weak to moderately developed slaty cleavage. Clasts in the broken formation are elongate, defining a NE-trending stretching lineation (mean orientation of 064°), and some of the clasts are asymmetric, with shapes suggesting a top-NE sense of shear (Fig. 9D). The kinematics of the deformation in this zone appear incompatible with the overall NW-vergence of the folds and cleavage, and we suggest it represents a localized zone of ductile shear post-dating the main folding, but predating the brittle normal faulting that follows (see below). The foliation in these rocks is overgrown by lawsonite, so that it seems likely that it formed during the subduction process, perhaps as a zone of backthrusting.

The Slab section as a whole is intensively disrupted by normal faults (Fig. 8). The faults dip gently either N or SE, and locally form sets soling onto near horizontal slip surfaces (Figs. 11 and 12). The mean orientation of the measured faults is 15/112. Kinematic data are not easy to find and vary widely in orientation, but the mean of slip lineations we have measured is NE, around 055°.

Thomes Creek: Lanz Section

The Lanz section (Fig. 13) is named after the Lanz pack-trail (no longer passable), which crosses Thomes Creek near the eastern boundary of the section. The western boundary of the Lanz section in Thomes Creek is very poorly exposed over 1 km, and the heavily degraded material seen in the canyon walls suggests the presence of one or more major faults. A discontinuous body of metavolcanic rocks ∼200 m across occupies the center of this zone of faulting. A simple interpretation is that the metavolcanics mark the stratigraphic base of the metasedimentary sequence to the east, and that they were disrupted by a major accretion-related thrust that placed the Lanz section above the Slab section to its west.

The eastern boundary of the Lanz section with the South Fork Mountain Schist (SFMS) is a discrete planar fault dipping 56/025, perfectly exposed on a strath terrace on the south side of Thomes Creek (Figs. 13 and 15). The boundary was named the Log Spring fault by Suppe (1973), who traced it north into the Tomhead Mountain area. Our location for the boundary in Thomes Creek is 700 m east of that shown by Suppe (1973), and 1300 m west of the contact shown by on the Willows geological sheet (Blake et al., 1992). The exposed contact separates very similar protoliths: dark graphitic pelite with dispersed fragments of greywacke sandstone, typical of the broken formation facies we have described from the Middle Eel and Slab sections. The rocks on either side are readily distinguished, however, most obviously by the strongly differentiated character of the primary foliation in the SFMS, which is absent in the Lanz section.

For several hundred meters west of the contact the stream bed is filled with boulders up to several tens of meters across of massive mafic blueschist. Just east of the contact, several blocks have selvedges of pelitic schist and metachert that are clearly derived from the SFMS. This blueschist body is not exposed in the stream, but the blocks appear to be derived from the bluffs on either side of the canyon. We suggest that a layer ∼100 m thick of mafic schist lies at or near the structural base of the SFMS, and that the fault exposed in the creek is in fact a late normal fault that cuts out this layer and the true Log Spring fault (Fig. 15). We assume that the Log Spring fault itself is a major thrust fault formed during underplating of the Lanz section beneath the SFMS.

The Lanz section is made up largely of thin-bedded metagreywacke sandstone and shale, some metaconglomerate, broken formation, and one or more bodies of metabasalt apparently floating as large blocks, or knockers, in broken formation (Figs. 12C and 12D). The metabasalt commonly shows well-developed pillow structure; it is moderately deformed, and forms slabs that thin out at either end. It is locally cut by narrow shear zones in which it has been converted to fine-grained glaucophane schist (Fig. 14A), but the bulk of the rock consists of relict igneous plagioclase laths in a matrix of fine-grained chlorite and sphene.

The metasediments show a pervasive foliation, phyllitic in character, but detrital grains are still preserved in the metasandstones (Fig. 14B). Lawsonite tablets are abundant in the pelitic rocks, and lie parallel to the foliation, suggesting that the fabric formed during or after the high-P/low-T metamorphism. Metaconglomerate and broken formation are strongly deformed, so that the clasts define a shape fabric (foliation and stretching lineation). The foliation is for the most part parallel to bedding, and together with bedding has predominantly gentle easterly dips through the section, although orientations are very variable on the outcrop scale. The mean stretching lineation is 14/074. Thin-bedded greywacke sandstones show tight early folds that are likely related to the main foliation: these are strongly overprinted by later folds (Fig. 14C) and have very variable orientations, but predominantly easterly plunges.

The primary foliation, together with bedding, has been crenulated and folded through most of the section on scales ranging from mms to ∼50 m, defining a series of predominantly E- to NE-vergent secondary folds (Fig. 14A). These have a differentiated crenulation cleavage developed parallel to their axial planes, but the new fabric is only locally strong enough to become the dominant foliation (Fig. 14C). The mean orientation of the secondary fold hinges is 09/135 (Fig. 4H), but they become very widely dispersed on the overturned limbs of major folds. Fold axial planes, and the crenulation cleavage, are regionally subhorizontal (mean 15/232), but in the eastern part of the section commonly dip gently east, less steeply than the main foliation, consistent with the easterly vergence of the folds (Fig. 13). The larger folds produce distinctive steep zones where bedding and the early foliation are around vertical, and in the eastern part of the section they locally have overturned limbs with W-vergent minor folds. The variation in orientation of the cleavage with strain (Fig. 14A), and the more intense thinning and disruption of bedding on the overturned limbs of the E-vergent folds, suggest that this deformation event was associated with E-directed shear.

In the western half of the section, secondary folds are much more variable in orientation and vergence. The variability may partly be a result of superposition on the earlier set of folds: interference structures are common in several parts of the section (Figs. 13 and 14C).

Within 500 m of the boundary with the Log Spring fault, the secondary folding becomes more intense, with overturned limbs on the major folds, and the crenulation cleavage locally becomes the dominant foliation visible in the field. This suggests that the deformation may be related in some way to the fault, a point discussed further below.

Thomes Creek: South Fork Mountain Schist

The South Fork Mountain Schist (SFMS) section is the structurally highest and easternmost tectonic unit in the Eastern Belt of the Franciscan. In the Thomes Creek transect it comprises two or more internally coherent thrust sheets discussed further below. Metamorphic grade in the Franciscan increases from westward, and the SFMS is the highest grade unit in this transect. It also shows the highest intensity of ductile deformation. The section through the SFMS extends from its lower boundary at the Log Spring thrust (described above) for 8 km downstream to its upper boundary, 1.2 km west of the Thomes Creek Gorge near Paskenta, giving a total structural thickness of ∼3.5 km. Due to the variable strike of the stream over this distance, the section has been projected onto multiple section lines (Fig. 15). Along its upper boundary at this point it lies in sharp contact with metasedimentary and volcanic rocks, attributed tentatively to the Galice Formation of the Klamath Mountains. The structurally higher portion of the Galice slice is coarse greywacke while the lower portion is a silty shale. The entire slice is less deformed and lower grade than the adjacent SFMS, as evidenced by a lesser degree of recrystallization and differentiation. The contact between the Galice and the SFMS appears to be parallel to the foliation in both rock units, and dips ∼60° ENE. We did not find any clear evidence for the nature of this contact, which must be a fault within the Coast Range fault system. Serpentinite in the overlying Tehama-Colusa mélange, however, does contain abundant shear surfaces, which provide some insight as to the nature of the Coast Range fault.

Pervasively developed reverse-sense shears exposed in the M4 forest road south of Thomes Creek dip moderately E, and slip lineations plunge toward 073°, giving a shear sense directed toward 253° (Fig. 18G); later more localized normal-sense shears dip NE, with slip lineations directed toward 023° (Fig. 18H).

The SFMS includes pelitic mica-schists, massive metagreywacke, minor metachert, the ∼900 m thick Chinquapin metabasalt member near the structural top of the unit, and thinner slices of metabasalt within and at the base of the unit. A few meters of broken formation, composed of sandstone rafts in a pelitic matrix, are present close to the Log Spring fault near the structural base of the unit (Fig. 16A). Massive thick-bedded greywacke units have been distinguished on the map and section (Fig. 15), but in many places pelitic schist and greywacke are interlayered on a variety of scales. Proportions of greywacke to pelite were not rigorously defined in this study and these interbedded sequences have been mapped as pelite. The metasedimentary units include quartz, white mica, albite, chlorite, lawsonite, calcium carbonate, and framboidal pyrite. Metabasalts include glaucophane, stilpnomelane, and Fe-rich epidote. Jadeitic pyroxene has not been observed in the SFMS in this study or by previous workers. Brown and Ghent (1983) report that pyroxenes from metabasalts in the Ball Rock and Black Butte areas are rich in acmite (50%–75%) and poor in jadeite (10%–30%).

There are four recorded episodes of deformation. The earliest detectable foliation, S1, consists of 1–2-mm-thick differentiated bands of quartz and white mica. It is preserved predominately within crenulation arcs at a high angle to the main foliation, S2 (Fig. 17A). S1 is preserved at the outcrop scale in limited locations within the westernmost massive greywacke unit and in just a single location east of the greywacke, where a green chert layer contains D2 folds whose axial planes can be seen to form the main S2 foliation of the surrounding rock (Fig. 17B). Although S1 is the result of the oldest discernible deformation episode, it may not represent the first deformation event experienced by the SFMS. It is possible that the deformation of the SFMS was intense enough to erase evidence of earlier episodes.

S2 is the main fabric and forms the foliation visible at outcrop scales. It is primarily visible as differentiated quartz and mica domains. Complete transposition has left S2 parallel to bedding in the SFMS. Lawsonite is commonly parallel to S2 and is particularly abundant in the mica-rich domains parallel to S2 (Fig. 17C). Where lawsonite and the foliation are at high angles to each other, S2 is observed to wrap around lawsonite grains, indicating that S2 was formed after the SFMS had reached blueschist facies conditions (Fig. 17D). The westernmost greywacke unit within the SFMS (Fig. 15) contains multiple outcrops with preserved tight to isoclinal D2 folds (Fig. 16C). Hinges are rarely exposed and are difficult to measure, but appear to be at an orientation which does not match that of folds in nearby pelitic material. Fold axial planes are commonly subparallel to the transposed bedding. Also present in this massive greywacke section are thin interlayers of pelitic material which host both D2 and D3 structures. D2 structures are expressed as mm and cm crenulations at a high angle to, and clearly refolded by, cm scale D3 crenulations (Fig. 20B). Stretching lineations on S2 are defined by deformed quartz granules and pebbles in coarse-grained metagreywacke (Fig. 16D), and deformed rock fragments in broken formation. Amphibole mineral lineations in metabasalt are parallel to the stretching lineation, and are likely to have the same significance. Taken together, these lineations have an average orientation of 49/013 (Fig. 18E). Kinematic indicators in the SFMS are sparse and largely restricted to asymmetrically boudinaged quartz veins. These are commonly ambiguous or give conflicting shear senses. Three S2 stretching lineations were also measured in Cottonwood Creek. They have an average orientation of 30/302, significantly more westerly trending than the lineations in Thomes Creek (Fig. 18E).

The third episode of deformation folded S2 into kink bands, crenulations, and asymmetric folds, which are pervasive throughout the SFMS (Fig. 16B). D3 folds are most strongly expressed in pelitic units and are weakest in sections which are dominated by metagreywacke or metabasalt. With the exception of an ∼30-m-long section observed in Cottonwood Creek, the D3 folds are generally absent from the metabasalts. D3 fold hinges in the Thomes Creek section of the SFMS primarily plunge toward the northeast, with an average orientation of 32/050 (Fig. 18A). Although the fold hinges have a preferred orientation, there is a fair amount of scatter and the trend varies by ± 35 degrees while the plunge varies by ± 30 degrees. In places where the hinges plunge at a shallow angle they often break the horizontal, resulting in hinges which plunge in opposite directions despite having similar trends.

Stretching lineations in the SFMS and D3 fold hinges have similar average orientations. If this were due to the folds being rotated into the extension direction then W- and E-vergent folds would have different average orientations. They are seen to have the same average orientation, indicating that they were not rotated (Fig. 18F), and we interpret the stretching lineations as being a result of D2 deformation.

The vergence of the asymmetric D3 folds is different on the opposing limbs of large-scale folds (Figs. 19A, 19B, 14E, and 14F), allowing for the identification of asymmetric folds up to hundreds of meters in amplitude. The large-scale folds are interpreted to have the same kink band or asymmetric fold structure as the smaller folds that are visible at outcrop scale. The long limbs of these large-scale folds tend to have gently dipping foliation and NW-vergent minor folds while the short limbs have steeper, E-dipping foliation, E-vergent folds, and more intense deformation.

In the extreme easternmost portion of the SFMS D3 folds are both more intense and more chaotic. The western end of the transect is dominated by the 0.7-km-thick steep limb of a major, SW-plunging D3 fold. Straddling the location where the steep limb breaks the vertical and becomes overturned are outcrops with fold interference patterns. On the upright portion of the limb, W-vergent, SW-plunging D3 folds with 15 cm amplitudes are seen to cross E-vergent and NE-plunging folds with 5 cm amplitudes. On the overturned portion of the limb, the vergences appear to be reversed due to the inversion of the layering (Fig. 20D), but the two phases of folding are readily identified by the differences in their orientations and amplitudes (Fig. 20A). Timing relationships between the two sets of folds are not clear. West of this transition zone, within the overturned limb, is another outcrop with two phases of folding. The dominant phase is defined by cm to m scale, symmetric to E-vergent D3 folds. A smaller cm-scale set of W-vergent folds is also present and may predate the D3 folds (Fig. 19D). Portions of the overturned limb contain a measurable, gently dipping S3. Within the dominantly E-vergent, overturned limb is a small, W-vergent, and apparently upright outcrop. Folds at the transition from the E-vergent section to the W-vergent section record opposing vergences on opposite limbs (Fig. 19A). Axial planes throughout this section are gently dipping in a manner consistent with the measurable S3 present at nearby outcrops (Fig. 19B). In places, S3 can be seen to fan out around D3 folds, contributing some variability to its orientation (Fig. 19C).

D3 fold hinges in Cottonwood Creek have an average orientation of 43/349 (Fig. 18B), ∼70° counterclockwise relative to the hinges in Thomes Creek. In both Thomes Creek and Cottonwood Creek, D3 fold hinges trend ∼40° clockwise relative to the stretching lineations in the same transect. Additionally, the dip direction of the main S2 foliation in Cottonwood Creek is oriented ∼50° counterclockwise from the dip direction of the S2 foliation in Thomes Creek (Figs. 18C and 18D). This suggests the possibility that the outcrop scale structure of the two sections was formed simultaneously and that a later, as yet unobserved, structure has rotated the Cottonwood Creek section around a vertical axis ∼50–70° counterclockwise relative to the Thomes Creek section.

The fourth and most recent episode of folding affects only limited portions of the SFMS, locally producing fold interference patterns (Fig. 20C) and folds with undulatory hinges.

A large fault with predominately reverse sense kinematic indicators, but also some normal sense indicators, bisects the Chinquapin member (Figs. 21A and 21B). We interpret this as evidence of thrust faulting followed by more minor normal sense reactivation. It is likely that this is the continuation of the Tomhead fault described by Worrall (1981), a fault which bisects the Chinquapin in Cottonwood Creek. The portion of the fault exposed in Thomes Creek is oriented at 63/080 and is accompanied by a stretching lineation oriented at 36/016. A sequence of cherty metasediments (Fig. 21C) structurally above metabasalt was observed directly below the Tomhead fault, indicating that the two thrust slices making up the Chinquapin are upright and young to the east. Chert is also present adjacent to the structurally highest and eastern most limit of the Chinquapin. The eastern portions of both the Cottonwood Creek and the Thomes Creek transects include faults with a reverse sense of motion, which juxtapose pelitic schist with metabasalt (Fig. 21D). Faults tend to dip E to NE, though there is some variability in Cottonwood Creek exposures, where both normal and thrust faults are present. At the western end of the SFMS, a thrust fault with a markedly different orientation of 52/208 truncates the hinge of a major D3 fold. Extensive fault parallel quartz veining is present, in places up to 25 cm thick. Based on the offset of the fold hinge, the fault appears to have accommodated a somewhat minor amount of motion. Minor normal faults are present throughout the SFMS; kinematic data are difficult to obtain, but the sense of motion can be recognized by Riedel shears and by fault drag folding.

Veining and pressure solution at a variety of scales are common throughout the SFMS. At the meter scale, younger veins can be seen cross cutting the foliation while older veins have been transposed and are parallel to the foliation. At the thin section scale, quartz has grown in pressure shadows around pyrite (Fig. 22A), in the dilational arcs of fold hinges (Fig. 22B), and in microcracks where the quartz has locally grown orthogonal to the walls. Pressure solution has differentiated the rock into quartz-rich and mica-rich domains and pressure solution seams truncate quartz grains, suggesting that precipitated quartz was locally sourced. Dilational microcracks are pervasive throughout the SFMS and are filled with quartz. Different quartz veins have experienced different degrees of dynamic recrystallization, likely reflecting different timing relative to ductile deformation (Figs. 22C and 22D). The coexistence of the microcracks and their varying degrees of ductile overprint indicates that the SFMS was experiencing brittle and ductile deformation simultaneously while under blueschist facies conditions, and at temperatures near the brittle-ductile transition.

DISCUSSION

Eastern Belt rocks in our transect contain a variety of structures that together record the stages of subduction, accretion, and exhumation. Deformation was likely continuous and it is not always possible to say with certainty which structures belong to which stage of the evolution of the accretionary complex, even when they can be chronologically ordered. Here we interpret structures that predate blueschist facies mineral growth to be related to subduction, and structures that are synchronous with blueschist facies metamorphism are likely related to accretion or the early stages of exhumation. Structures formed during the later stages of exhumation are identified by their predominantly brittle nature, relation to extension (vertical shortening), and lack of associated high-P/low-T minerals.

A potentially controversial issue is the age and origin of the disruptive deformation that produced the broken formation found throughout the region, including the westernmost (structurally lowest) part of the SFMS. Critical to this discussion are the following observations. (i) Clasts in the broken formation primarily comprise rafts or fragments of greywacke sandstone that have irregular edges and embayments. Fragments of metavolcanic rocks are uncommon, and we found no blocks with metamorphic grade higher than their surroundings. (ii) The degree of fragmentation in the broken formation appears to be unrelated to the intensity of the deformational fabric, and some bodies of broken formation have little or no fabric. (iii) Many of the bodies of broken formation appear to form stratigraphically bound interlayers in bedded greywacke sandstone. (iv) Layers of broken formation commonly contain elongate clasts oriented parallel to their boundaries, and the fabric in the mélange is folded in the same way as the surrounding bedded sequences of greywacke sandstone. These features are consistent with an origin of the broken formation by surficial sliding and mass transport in a trench environment, as suggested by Wakabayashi (2011) and Platt (2015).

Structures related to subduction are widely developed but commonly overprinted by later deformation. These structures include early thrust repetitions of basalt-chert-greywacke sequences in the TMC. Limited kinematic data from these structures suggest roughly NE-directed shear, which is not obviously consistent with the likely direction of thrusting in the Franciscan accretionary wedge. They may have formed by backthrusting during subduction, but given the complexity of the later deformation, and the likelihood that convergence was oblique during at least part of Franciscan history, vertical-axis rotation of early formed structures can be expected (see, for example, Platt, 2000). Vertical-axis rotation is also suggested by the systematic differences in orientation in the SFMS between the Thomes Creek and Cottonwood transects.

N- to NW-vergent asymmetric folds accompanied by a slaty cleavage are common in the lawsonite-albite facies rocks of the Middle Eel and Slab transects; the cleavage is overgrown by lawsonite, suggesting that these structures are related to subduction. A NE-directed zone of ductile shear a few tens of meters wide in broken formation in the Slab section also predates lawsonite growth, and may be a backthrust formed during subduction. The earliest foliations found in all parts of the transect are likely related to subduction. Sheet vein complexes parallel to early foliation in the TMC carry jadeite and lawsonite; the first cleavages in lawsonite-albite facies rocks of the Middle Eel and Slab sections are overgrown by lawsonite; the first cleavage in the Lanz section contains oriented stable lawsonite and is overgrown by lawsonite porphyroblasts. Early folds in the Lanz section are coeval with the first foliation, but we were not able to determine their vergence. The earliest deformational fabric in the SFMS likely predates lawsonite growth, but is heavily overprinted by later deformation.

Contractional structures that formed during the high-P/low-T metamorphism represent deformation during subduction or accretion at depth (underplating). These likely include the fault boundaries between the various large-scale tectonic units in the region. The TMC, for example, is in thrust contact with underlying lawsonite-albite facies rocks, and was likely emplaced under those conditions, as the tectonic boundaries are marked by sheeted vein complexes containing lawsonite. The original Log Spring thrust, and the likely thrust contact between the Lanz and Slab sections, are not exposed in Thomes Creek, but we think they are likely to have formed during continued subduction-related convergence. The large-scale synform in the TMC postdates subduction related thrusts, but predates exhumation-related normal faulting, and may have formed during the accretion stage.

The second-generation folds in the Lanz section, and both D2 and D3 in the SFMS, all demonstrably formed under high-P/low-T conditions. Secondary folds and associated crenulation cleavage in the Lanz section deform a lawsonite-bearing fabric (S1), but formed during glaucophane growth in metabasaltic rocks. Lawsonite in the SFMS formed before or during the main foliation (S2), and lawsonite porphyroblasts locally show rotational inclusions suggesting continued growth during D2. Hence these structures all formed after the initial phase of subduction, but are likely to represent continued subduction-related shear during or after accretion. The same is true of the Log Spring fault, which is clearly a contractional fault, but which juxtaposes rocks with significantly different metamorphic and deformational histories. It is therefore likely to represent the subduction zone interface during underplating of the Lanz section under blueschist facies conditions. This therefore raises the question of its relationship to the deformational history of the SFMS and Lanz sections on either side of the fault.

The E-vergent secondary folds that are dominant in the eastern part of the Lanz section increase in intensity toward the Log Spring fault, and they have the same orientation and style as E-vergent folds in the adjacent western part of the SFMS. The E-vergent folds in the Lanz section are related to E-directed shear, however, whereas those in the SFMS formed on overturned limbs of the major W-vergent D3 folds that are predominant throughout the SFMS, and are likely to be related to W-directed shear. A possible solution to these paradoxical relationships is that the E-vergent folding in the Lanz section was related to E-directed backthrusting during the initial stages of underplating, and that these structures were crosscut by the W-directed Log Spring fault, which accomplished the final emplacement of the Lanz section beneath the SFMS. D3 in the SFMS likely accompanied motion on the Log Spring fault.

The intensity and complexity of the deformation in the SFMS also increases on the east side of the section, and includes several syn-metamorphic thrusts, including a major thrust that duplicates the mafic section in the Chinquapin metabasalt. This suggests that subduction-related shear continued along the eastern tectonic boundary of the Franciscan Complex as a whole throughout the underplating history.

Normal faults related to exhumation of the high-P/low-T rocks are widespread and locally very abundant. Normal faults in the TMC cut structures related to subduction and accretion, and are associated with sheeted vein complexes that carry lower pressure assemblages (lawsonite + albite) than those formed at peak pressure (lawsonite + jadeite). The precise timing and geometrical relationships between these normal faults and tectonic boundaries between the TMC and adjacent lawsonite-albite facies rocks are unclear; they appear to have formed under comparable conditions.

Normal faults are particularly abundant in the Slab section, and dominate the structure. They are less obvious in the Lanz and SFMS sections, but we observed a probable normal fault with a displacement of ∼100 m that cuts out the Log Spring thrust in Thomes Creek, and normal sense reactivation of the synmetamorphic thrust in the Chinquapin metabasalts of the SFMS. Large-scale normal-sense shear is likely to have occurred along the Coast Range fault, which played a major role at a late stage in the exhumation of the Franciscan Complex as a whole, and we have documented both reverse and normal sense motion in sheared serpentinite from the Tehama-Colusa mélange.

The presence of quartz veining in structures related to all phases of evolution, and the differentiated character of all the foliations throughout the profile, indicate that pressure solution was a significant deformation mechanism throughout the entire evolution of the Eastern Belt and highlights its importance in facilitating the deformation of this part of the Franciscan Complex.

CONCLUSIONS

The Eastern Belt of the Franciscan Complex includes rocks metamorphosed under a variety of high-P/low-T conditions, and progressively deformed during a history of subduction, underplating, and exhumation. The earliest structures are stratigraphically bound layers of broken formation, and isoclinal folds in sandstone that predate the main cleavage; these features are likely related to surficial processes in the trench environment. Subduction-related structures include imbricate thrusting in the blueschist-facies Taliaferro Metamorphic Complex, early W-directed folds and cleavage in lawsonite-albite facies metagreywackes, and the earliest foliation in the blueschist facies South Fork Mountain Schist. Structures formed during underplating include E-directed shear zones and E-vergent folding related to backthrusting within the accretionary wedge; major W-directed syn- to post-metamorphic thrusts that juxtapose tectonic units of different metamorphic grade; and intensive W-vergent folding and crenulation cleavage in the South Fork Mountain Schist, which are likely to be associated with the major syn- to post-metamorphic thrust faults bounding the unit, including the Coast Range fault system on the eastern margin of the Franciscan. Structures related to exhumation include locally intensive normal faulting throughout the section, and normal-sense reactivation of the Coast Range fault.

ACKNOWLEDGMENTS

This research was funded in part by National Science Foundation grant EAR-1250128 to J. Platt. We are grateful to Whitney Behr, Alex Lusk, Ellen Platzman, Daniel Platt, Daniel Schmidt, and Francisco MeldeFontenay for their help in the field, and to Tom MacKinnon for sharing the results of his work in Grindstone Creek and for numerous stimulating discussions. We appreciate constructive and helpful reviews by Gary Ernst and John Wakabayashi, and we thank Damian Nance for editorial handling.

Gold Open Access: This paper is published under the terms of the CC-BY-NC license.