Figure 1. View looking SW at Shenandoah Valley from Miller’s Head (3484 feet elevation). The forested hills to the center-left are Precambrian metamorphic rocks (1600 – 1000 Ma), separated by an arcuate fault from Paleozoic sedimentary rocks (540 – 250 Ma) that get younger to the west. The ridge in the distance is constructed of Devonian rocks (416 – 360 Ma) whereas the closer ridge is Ordovician to Silurian sandstone (443 – 419 Ma). Today’s post will examine the Precambrian rocks that comprise the Blue Ridge Mountains, which we saw in a the last post and in a previous post.

Figure 2. (A) Map of northern Virginia (NOVA) showing my home (star), Washington DC, and the study area in Shenandoah National Park (rectangle in lower left). The inset photos show the stunning views to be found in the lower part of the study area. (B) Geologic map of the study area, showing Precambrian rocks in pink shades and Cambrian rocks in brown, Ordovician in green, and Devonian strata in orange shades. The Shenandoah Valley forms a syncline with smaller folds contained within it, like folds in a rug. Younger rocks are exposed by erosion along the axis of a syncline. The solid lines are faults that have been identified in the field although the kind of fault can’t always be determined.

Figure 3. Images from Hawksbill Mountain (elevation 4042 feet), the highest point in Shenandoah National Park. (A) View looking north, showing bedding planes dipping to the SE as we’ve seen throughout NOVA. This is the general structural trend along the eastern margin of North America. These volcaniclastic deposits are part of the Catoctin Formation, dated by radioisotopes to between 1000 and 485 Ma. (B) A close-up shows that these rocks are fissile, which means they are forming thin layers as they weather. They were originally very fine grained, possibly ash or mud (clay minerals) and other weathering products. (C) This photo reveals (despite the dappled shade of trees) a contact (yellow line) between tan and bluish rock that has no other distinguishing features. This must be caused by a slight difference in composition and/or texture that leads to subtle variations in reflected light; if I may speculate, I think the blue ash/sediment filled a depression in the tan material; however, this conjecture is based on the physical environment when this volcanic material was produced. Eruptions were not continuous and there was always a slightly weathered surface upon which new ash was deposited. (D) Outcrop of Catoctin Formation rocks about 600 feet lower than the summit, and much further down-section from the rocks seen in plate A. This basalt/ash was deposited millions of years before what we see in plate A and the magma chamber would have evolved substantially. The original bedding is, as near as I could tell, nearly horizontal rather than tilted to the SE. This implies that brittle deformation (i.e. faulting), which occurred tens of millions of years after eruption, subsequent burial, and metamorphosis, was localized within the larger body of Catoctin Formation rocks (3000 feet thick and extending for many miles); in other words, these metabasalts were hard as rocks (as they say) when they were compressed horizontally. Perhaps they were even transported tens of miles along a thrust fault? We saw evidence of a Precambrian thrust fault at Bull Run Nature Preserve.

Figure 4. Catoctin Formation metabasalts at Dark Hollow Falls. (A) These rocks don’t form sheer cliffs, so this creek flows over a series of ledges for a vertical distance of about 70 feet.(B) The bedding is tilting to the SE as at Hawksbill summit and contains fissile layers as seen here, intercalated with massive units as seen in plate A. Also note the dark layer in the center of the photo, which may be similar in origin to the blue layer seen in Fig. 3C. (C) Representative boulder of the rocks at Hollow Falls. Notice the white flecks in this sample, which is about 18 inches in length. (D) Close-up of the larger light-colored ellipse in the lower-left part of plate C, showing concentric rings of light material (probably quartz) in a fine matrix (probably basalt). This is an amygdule, which is a mineral filling a cavity in a volcanic rock that is filled with vesicles. The vesicles are originally pockets of volcanic gasses that are trapped when the basalt is erupted, and then fill with hydrothermal fluids which deposit minerals. (E) Bedding surface showing many semicircular ridges, which are (probably) remnant from when this basalt was exposed to the atmosphere and the gases escaped; in other words, this was the top of a basalt flow whereas plate D was too deep within the flow for the gas to escape before the basalt solidified. A moment in time frozen for more than 500 million years.

Figure 5. Photos of Miller’s Head. (A) View looking west (see Fig. 2B for the geologic map). The yellow lines approximately outline faults within the Paleozoic rocks underlying Shenandoah Valley; note that one curves to the west where Neoproterozoic (1000 – 542 Ma) rocks jut into the valley. The second fault borders a low ridge comprised of the same rocks we found on this peak. (B) Vertical joint that shows intersecting joints (the X’s seen in the rock face). Water seeps in through this system of fractures and weathers the rock into blocks. (C) The result of this weathering process is a mountain covered by a veneer of rubble. The entire mountain in this area is turning into a pile of boulders that look like they were pushed aside by a bulldozer, but they haven’t moved other than sliding over one another down the steep slope. (D) Close-up (4X magnification) of a fresh surface of this rock, which is Charnockite, a granitoid rock that contains pyroxene minerals, which do not occur in granite. This sample reveals quartz (Q), plagioclase feldspar (Pf), potassium feldspar (Kf), and a lot of pyroxene (Px). Note that Pf is white and not dark, which would be more indicative of a mantle source, and pink Kf which is typical for a granite formed from crustal material. Charnockites are enigmatic and almost entirely found in Precambrian rocks. According to RockD (radiometric dating), these intrusive rocks are between 1600 and 1000 million-years old.

Summary. This post doesn’t add anything new to what we’ve already learned from previous field trips but it reinforces the picture that has been developing from our previous posts in NOVA and elsewhere; tectonic plates were colliding along the eastern margin of North America as long ago as 1.6 billion years, while muddy sediments were being deposited in deep water (below wave base or rivers), and continued doing so until about 500 million-years ago. This was a discontinuous process and, considering the billion years duration of this tectonic upheaval, it is possible that multiple mantle plumes were competing for space. Such a huge span of time could easily encompass more than one Wilson Cycle, but the best I can say in this post is that the Proterozoic (2500 – 542 Ma) looks to have been as active an age as we’ve seen in the last 500 million years.

One final note. The earth is cooling very slowly but, nevertheless, it was hotter in the Precambrian. This means that plate tectonics, driven by mantle upwelling (i.e. plumes) would have been more vigorous although not by an order of magnitude. Thus, given the immense span of time between the Middle Proterozoic (1600 – 1000 Ma) Charnockite we encountered (Fig. 5) on this trip and the Catoctin Metabasalts (1000 – 485 Ma), it is safe to say that we haven’t seen the whole story.

The rocks speak softly but they know the truth …

Figure 1. View looking west from Dickey Ridge/Hill (See Fig. 2) towards the Shenandoah Valley. This is where the Ridge and Valley province begins, a series of elongate mountains running approximately 30 degrees east of north, the same as the structural trend we’ve seen in previous locations throughout northern Virginia. Dickey Hill has an elevation of 2427 feet, about 1800 feet higher than the valley floor.

Figure 2. The left image is a map of Shenandoah National Park, which extends along the ridge line of the Blue Ridge Mountains. The study area is circled and a geologic map from RockD is shown in the right image. The blue dot is where we started our climb to Dickey Hill. The rocks of this area (shown in a light-gray color) are the Catoctin Formation (1000 – 485 Ma): metabasalt (metamorphosed to greenschist facies), including some preserved volcanic structures, which are the subject of this post.

Figure 3. Exposure of Catoctin metabasalts, showing original bedding, which is dipping away from the camera, and variability of these rocks in outcrop.

Figure 4. Example of what metabasalts look like in the field. (A) boulder showing a greenish hue that has been roughed up a bit and is no longer angular. (B) Close up of area enclosed in (A), showing some of the biological materials that produce the mottled appearance of these rocks. The bright areas with angular form are probably quartz and feldspar, two minerals commonly associated with greenschist facies metamorphism (low pressure and low temperature).

Figure 5. Example of textures found in Catoctin metabasalts: (A) bluish-green boulder with angular form. (B) Close up of fine-scale textures preserved from the original basalt. The acicular sections result from a preferred alignment of minerals, indicating a flow direction; blocky texture indicates that this part of the sample was part of a larger flow and probably more viscous; nodular textures probably result from weathering at some point, possibly soon after deposition. Volcanic flows are not uniform; for example, the basalts of the Catoctin formation are 3000 feet thick but were extruded over tens of millions of years, during which the magma chamber would have changed in composition. These “nodules” are elongate and could be due to recent (last few million years) weathering of acicular textures. Perhaps nodular is the wrong word; these also look a lot like miniature pahoehoe lava, which is ropy when first created.

Figure 6. (A) angular boulder with a rich blue color, suggestive of blueschist facies metamorphism (low temperature and high pressure). (B) Close-up showing ropy, flowing structures around fragments of basalt (originally) that were carried with the flow. We saw this kind of flow structure at Koko Crater on Oahu in a previous post, but those rocks were only a few million years old and hadn’t been buried and subjected to huge tectonic stresses.

Figure 7. Images of volcanic textures in an outcrops. (A) This photo shows a slight difference in texture between the intrusion and the surrounding lava, even though both were erupted at the same time. The lava produced by a magma chamber, especially near the surface, is not homogeneous but rather a poorly mixed assortment of molten and solid material infused with high-pressure volcanic gas. The circled area labeled as a Cavity (speculative) is an example of this heterogeneous volcanic texture. (B) This photo of an outcrop shows angular fragments quite distinct from the background matrix (not to be confused with lichen); these are not original textures because basalt does not contain light-colored rock fragments. These inclusions are metamorphic in origin, probably quartz and feldspar.

Figure 8. This hand sample was photographed at the top of Dickey Hill. It is as fresh a sample as you can get without a rock-hammer. Note the thin filament of material separating two conchoidal fracture zones. The greenish color is why metabasalts are called greenschist.

Figure 9. This photo shows a couple of post-tectonic textures that reflect events after these volcanic rocks were buried and metamorphosed. Joints are brittle fractures that occur when a rock has been exhumed and the stress regime has reduced; the rocks break in patterns like the “X” that has been superimposed on this image. Joints cannot be dated so all I can say is that this pattern, which was expressed in all of the outcrops I saw, occurred millions of years after burial, probably after the break-up of Pangea, when the mountains that once overlay this area were eroded away. The circled area shows rounded corners like we saw in Fig. 5B, suggesting that water flowed over this outcrop for a long period of time. The fracturing in Fig. 8 would have occurred during this period of relaxing stress.

SUMMARY. About one-billion years ago, lava flowed onto the land that later became Virginia for millions of years, culminating in a 3000-foot-thick pile of basalt. This is half as thick as the Deccan Traps basalt province in India or the Columbia Plateau of North America; the latter was produced in 10-15 million years. Both of these geologic provinces are associated with collisional tectonic regimes.

The uncertainty in age for the Catoctin Formation (1000-485 Ma) is due to the uncertainties of radiometric dating, caused mostly by loss of radioactive products over time (loss of products gives false young ages); thus, it is probably safe to say that these rocks were originally produced about one-billion years ago.

We saw these rocks at Morven Park and Catoctin Creek, and we saw contemporaneous sedimentary rocks at Bull Run nature preserve. If these basalts are analogous (that’s a big IF) to the Deccan Traps and the Columbia Plateau, this was a collision of tectonic plates. This tentative interpretation is supported by the lack of pillow lava structures (lava erupted into deep water) reported in the Catoctin Formation. The presence of terrigenous sedimentary rocks deposited 500 million years later in Virginia suggests that an entire Wilson Cycle (opening and closing of an oceianbasin) occurred between the Catoctin formation (~1000 Ma) and the Harpers formation (~538 Ma).

The greenschist metabasalts of the Catoctin formation weren’t buried deeply or heated very much, so they weren’t close to the point of impact during the closing of Iapetus Ocean, which was somewhere on the continental shelf a hundred miles or more east of the study area. This tectonic collision began north of Virginia about 550 million-years ago (i.e. when the Harpers formation was deposited). It is called the Taconic Orogeny. The result was the supercontinent Pangea, which lasted several-hundred million years before being torn apart about 200 million-years ago.

Tectonic plates are dancing but who’s playing the music?

Mantle plumes are elusive for humans to track. Imagine a pan of boiling water; bubbles appear, some large, some small, spaced at seemingly random distances apart. They last only as long as it takes them to rise to the surface.

The rocks are trying to tell us what dance they are moving to …

Figure 1. The Waterford mill was constructed in the 1820’s and operated until 1939. It was located along Catoctin Creek because there is a significant drop in stream elevation here. A low ridge behind the mill suggests part of the reason: resistant rocks that produce a series of sudden drops in the stream, perfect for collecting water behind a dam not too far from the mill. The shorter the length of the “head race”, the lower the construction costs. The “tail race” can be seen leading from the waterwheel. This small ditch leads around the facility and drains into Catoctin Creek a couple hundred yards downstream. This post will explore upstream on Catoctin Creek, as far as the mill pond (storage pond for hydraulic pressure), and see what the rocks tell us.

Figure 2. (A) Regional map of the study area. My house is labeled with a star. Bull Run fault is a continuous normal fault that we saw at Morven Park and Banshee Reeks Nature Preserve; its location in the figure is approximate. It demarcates younger Mesozoic rocks to the east from older, Precambrian, rocks to the west; the latter form an elongate ridge that cuts through the study area (rectangle centered on Waterford), as well as the Blue Ridge further west, which front the Shenandoah Valley. Vertical movement along Bull Run fault began about 220 Ma (Ma is a million years, but measured by radiometric techniques) when Pangea was torn apart by the newly forming Atlantic Ocean. (B) Geologic map from RockD of Waterford area. The section of Catoctin Creek discussed in this post is enclosed by a blue ellipse. Note that the Precambrian rocks fall into two distinct sequences: 1.6 Ga (billion years measured with radioisotopes) and gneiss, both metamorphosed; and 1 Ga to 600 Ma volcanics and associated sedimentary rocks. The contact between them is an unconformity but the type cannot be determined from field data. This might be a buried and disrupted late Precambrian (~600 my) thrust fault that pushed older rocks over younger, like we saw at Bull Run Nature Preserve.

Figure 3. (A) View looking downstream along Catoctin Creek, showing rounded cobbles (~2 inches) in a sandy matrix, with silt and mud. This unlikely assortment of sediment grain sizes suggests to me that there are multiple sources being mixed along the creek; for example, rounded cobbles suggest miles of transport along swift-flowing creeks whereas mud is the product of physical and chemical weathering of rocks with a small quartz content (quartz is very hard and chemically stable; i.e. sandy beaches). (B) Exposure of older Precambrian schist along the creek bed, forming a low obstruction. (C) The schist layers (schist is a fissile rock) present a weathered appearance; chemical weathering produces mud in-situ without bedload transport, producing few cobbles.

Figure 4. Examples of different rock types found along Catoctin Creek. (A) Collection of angular schist boulders (~one foot in size) at one of the tributaries to the main stream. (B) Quartz intrusion, probably from the oldest rocks (metamorphosed granites and gneisses). The sample is about one foot long. (C) Fresh surface of the schist, showing fissility and a sheen associated with lower-grade metamorphosis, such as in phyllite. This sample was several feet long and had been transported to the mill-pond dam during construction of the mill.

Figure 5. The dam constructed to retain water for Waterford mill contains a variety of large boulders, but the majority were schist (see Fig. 4C). I didn’t see any granite or gneiss, which isn’t surprising because these large stones would have been difficult to transport in the early nineteenth century. They used what was readily available; the exposure of schist along the creek bed (Fig. 3A) suggests that the ridge fronting the mill (Fig. 1) was a likely source of material; after all, rock had to be removed to build the mill and the town of Waterford.

SUMMARY.

More than 1.5 billion years ago, something was happening in Loudon County, Virginia, long before there were multicellular organisms (eukaryotes) or even land plants. There was a collision of tectonic plates massive enough to produce granite and gneiss (high-grade metamorphism) and then deform these very durable rocks.

Four-hundred-million years later, an episode of extreme volcanism occurred and thick sequences of basalt and volcaniclastic sediments were laid down in Loudon County at about the same time (give or take a hundred million years) as deeply buried shales were being transformed into schist at several locations along the eastern margin of modern North America: less than 50 miles from Waterford, at Great Falls; what would become New York City; and Vermont. This was the closing of Iapetus, which took hundreds of millions of years and stretched for thousands of miles, creating Pangea and the rise of eukaryotes, fish, amphibians, reptiles, and mammals, not to mention land plants.

About two-hundred-million years ago, Pangea was torn apart and grabens formed, filling with sediment eroded from the surrounding elevated terrain. These sedimentary rocks are found east of Bull Run fault (Fig. 2A) where they remained protected from the elements (chemical and physical weathering) while the older rocks were elevated to form mountain peaks in the modern world and eroded into boulders, sand, and mud.

Two-hundred-million years later, the shattered remnants of a once-majestic mountain range, stretching from Canada to the Gulf of Mexico, comprise its core of metamorphic and igneous rocks recording events we can only speculate about today.

We discussed the tumultuous history of Precambrian rocks in Montana in my last post. The story of crustal shortening in western North America continues to this day. The huge, shield volcanoes comprising the Cascades Mountains show that this westward motion has not ceased at the current time.

The story of oceanic subduction and collision with multiple microcontinents is recorded in the rocks I had to drive past, so I have to resort to a geological map again.

Summary. This was a short post because I was occupied and didn’t have time to explore this fascinating region. Nevertheless, I can confidently say that when Pangea broke up, the North American tectonic plate began to “swallow” the proto-Pacific plate and any microcontinents it harbored.

This 200 my long process created the Rocky Mountains, the overthrust belts of Montana, the Black Hills of South Dakota, the Colorado Plateau, the volcanic Cascade Mountains, the complex system of faults that define California, and so many other geological features of the western North American craton.

It wasn’t as if another gigantic mountain range could form in the aftermath of Pangea’s break-up. The earth can only produce one of them every couple hundred million years, a tectonic pattern called a Wilson Cycle.

This post is experimental and not particularly interesting, but it is the best I can do under the circumstances; I followed Interstate 90 through the Rocky Mountains at 70 mph, with no pull-offs, and trying to take photos in the heavy traffic would have been suicidal. Instead of including a map, photos of outcrops, and some close-ups to examine mineralogy, I am relying on geological maps and my memory. The most experimental part is that I’m working on an iPad, which is a blessing and a curse. Let’s see how it worked out.

Summary. The oldest rocks (Precambrian metasediments shown in purple shades) are scattered throughout the Rocky Mountains. These old rocks were pushed and pulled for hundreds of millions of years as microcontinents collided to form what we call western North America.

Paleozoic rocks (500-230 my) that would have been deposited on top of them, or intruded into them, are only found in scraps here and there (I’m speculating, but Paleozoic rocks have a habit of turning up in the unlikeliest places).

During the late Cretaceous (about 80 my), granitoid intrusions forced their way into these older rocks, as I saw at Butte and other small mountain ranges (pink and tan hues). This was a geologically active period in the evolution of western North America.

About fifty-million years ago, volcanoes formed along the western margin of North America (e.g. Mt Hood and other volcanoes produced thousands of feet of volcanic rock, forming the Columbia plateau (yellow shades in Fig. 5). At approximately the same time (50 – 5 my) sediment was collecting in lakes and shallow inland seas leftover from the Cretaceous Interior Seaway. These sediments are undeformed and not very well lithified (i.e. not buried deeply); they appear east of Bozeman MT as green in Fig. 5.

Hidden beneath the Precambrian rocks, which were pushed eastward as much as 150 miles in Canada, and Miocene sediments, lay the oil and coal rich Cretaceous sediments laid down between about 150 and 60 my ago. As proof of this, Billings MT (rightmost circled area in Fig. 2), with a population less than 150 thousand, has three oil refineries; but it is so remote, with so little infrastructure (e.g. pipelines), that trucks deliver refined petroleum products to rail cars. It is a modern western boom town.

We’ll see what tomorrow brings …

This post is a little weak but I wanted to show that we can find evidence of the earth’s history in our back yard (literally). My interpretation may be completely wrong but it is consistent with my observations and (limited) understanding of volcaniclastic deposits.

I’m going to look at some more curious volcaniclastic deposits tomorrow …

Summary. Understanding volcanic stratigraphy is easy with a simple exercise. Spread your left hand out on the table, fingers apart. Each finger is a volcaniclastic flow, either pyroclastic, lava, or a lahar, separated by hundreds of thousands of years. Now, lay your right hand over the left but not with the fingers aligned. Imagine doing this hundreds of times, while peeling away the tops of your fingers randomly (i.e. erosion).

Remember the violent eruption of Mt. St. Helens? It was a pyroclastic eruption (mostly red-hot ash) but what made it destructive was the boiling hot mud encasing boulders of older volcanic material. The blast flattened the trees for miles and the lahar cleaned up the debris.

Imagine such an eruption occurring every year … thank god they are separated by centuries or millennia.

There’s only so much energy available, even for the dynamic earth …

This post is being written in Portland, Oregon, 2800 miles from Northern Virginia, where this journey began. I’m working on an iPad, which is new to me, so I’m going to limit this to a summary of previous posts, and briefly discuss some rocks I haven’t discussed before. I’ll go into more detail on the return trip, which will, however, take a different path.

I have said a lot about the rocks in NOVA (Northern Virginia), so I’ll refer to those posts. The eastern end of Fig. 1 is underlain by rocks more than one billion years old that record a collision on a continental scale.

We also found evidence of deposition in marine and coastal settings throughout the Paleozoic (~500 to 250 my), which I’ve discussed before. This period of erosion was interrupted in the Triassic Period, about 200 my ago, when the east coast began to stretch; coarse sediments filled newly developing basins throughout NOVA.

Our westward journey took us through Maryland and Pennsylvania (see profile above), where we found evidence of broad, shallow seas to the west and rising highlands to the east throughout the Paleozoic.

West of the ancestral Appalachian mountains, from Ohio to Illinois, we saw rolling hills covered with farms that replaced primordial forests. I don’t have any photos of this area, but there isn’t much to see. However, this is where extensive glaciation becomes evident, continuing across the Great Plains to Nebraska. I wrote about the moraines that dominate this region in a previous post.

This post picks up the story in eastern Wyoming (see profile above), where we find sediments deposited in coastal areas during the Late Cretaceous (~100 my) dominating the region.

Figure 2. Geologic map of the area around Rawlins, Wyoming. The majority of the rocks (green hues) are Cretaceous (~145-65 my). Faults (black lines) separate these older nearshore sands and muds from Miocene (~20 my) coarse sediments (yellow colors), Paleozoic marine sediments (aqua tints), and Mesozoic nearshore sediments (blue hues). Note the arch form of the Mesozoic layers; this is an anticline, folded layers of rock, the result of crustal shortening (i.e. compression).

SUMMARY. We started out in NOVA, where a titanic collision occurred more than 500 million years ago. We saw evidence of a similar orogeny in the Precambrian rocks exposed in Utah, Wyoming, and Idaho, long before they were deformed and pushed eastward.

During the Paleozoic era (500 – 230 my), thick layers of sediments were deposited in Pennsylvania (see Fig. 1) as the ancestral Appalachian Mountains rose, then eroded over hundreds of millions of years. The proto-Atlantic Ocean (Iapetus) opened and closed during this immense span of time.

We saw similar Paleozoic sedimentary rocks in Utah and Idaho (no photos available) but the big picture of continuous deposition along huge swaths of what is today western North America is recorded elsewhere (e.g. the Grand Canyon and Colorado Plateau).

The Mesozoic era is mostly recorded in sediments associated with the break-up of Pangea in NOVA, where stretching of the crust produced ridges and intervening grabens filled with coarse sediment. The Mesozoic and Cenozoic eras were spent eroding the ancestral Appalachian mountains as Eurasia and North America went their separate ways.

Vast expanses of shallow marine and lacustrine sediments were deposited in the (modern) central and western United States during the Mesozoic, accompanied by the eastward push of older rocks by episodic collisions; this was not a continental collision but probably a series of micro continents and island arcs being absorbed as oceanic crust was subducted. A vast interior seaway reached from the Gulf of Mexico to the Arctic Circle at this time.

The Cenozoic saw the eruption of vast quantities of volcanic material in Oregon and Idaho as Mt. Hood (and other volcanic centers) reached its peak of activity. The Cretaceous seaway dried up and terrestrial sediments replaced marine deposits, as the Colorado Plateau rose more than 5000 feet, shedding sediments everywhere. Finally, great ice sheets carved the earth’s surface into a new landscape defined by moraines and glacial valleys.

The Cenozoic is mostly under represented in NOVA because sediment collected on the continental shelf of North America, which was (and still is) a passive margin. Everything that was carried westward by the Mississippi River system was deposited ultimately in the Gulf of Mexico, where huge oil and gas fields developed from organic material transported by an ancestral Mississippi River drainage system. There was no room for sediment as the Appalachian mountains rose, in response to the removal of miles of overlying rock.

This has been a brief and probably inaccurate comparison and contrast of eastern and western geology of the United States, but it is only what I’ve seen with my own eyes, enhanced by the vast knowledge accumulated by generations of geologists tying the story together.

We’ll see what I learn on the return journey …

Figure 1. This photo shows the post-glacial, Pleistocene surface of the Carboniferous (320-300 Ma) limestone we’ve seen throughout Southern Ireland. The paucity of stone walls in the area suggests that boulders deposited by a glacier were less common here; the absence of drumlins or moraines further suggests that this area was scraped clean of loose material before the glacier retreated for the last time about 20 thousand years ago.

Figure 2. These layers of limestone are dipping at less than 20 degrees (field guesstimate), revealing different strata within the section. Low hills within the area indicated by a pin in the inset map of Fig. 1 were constructed of such tilted strata.

Figure 3. Geologic map of Ireland with the region discussed in this post highlighted by a black rectangle. The blue areas are the limestones we’ve seen before, indicated in the stratigraphic column (right of figure) with a matching arrow. These are the youngest rocks in this area. The pink area at the top of the study area is Ordovician (gray arrow in stratigraphic column), between 495 and 440 million-years old; these sedimentary and volcanic rocks are much older than the limestone. We’ll be focusing on Cambrian (545-495 Ma) sandstones that have been metamorphosed after burial and heating. This is the yellow-shaded area on the map and stratigraphic column. These are the oldest rocks we’ve seen in Ireland. Note the red area, and arrow in the lower map key (igneous rocks are shown separately in this map). These are Ordovician in age and thus much younger than the metasediments we’ll be looking at in this post.

Figure 4. The inset map focuses on the upper-left portion of the inset map in Fig. 3. The three main rock types introduced above are labeled for clarity; note the presence of older (Cambrian) metasediments sandwiched between intrusive granite and sedimentary and volcanic rocks of Ordovician age. This curious stratigraphic relationship is due to an orogeny that occurred during the Ordovician geologic period. Cambrian sediments, mostly sand and shale, were buried several miles beneath the surface, when even deeper rocks melted to form magma, which rose as granite. The sediments were heated under moderate pressure, and then injected with fluids from the magma. The Ordovician sediments were deposited during this orogeny but were not buried as deeply as the subjacent Cambrian rocks.

Plate A shows a typical exposure of the Cambrian metasediments. Note the general appearance of a bedding plane, tilted steeply, facing towards the camera. Plate B is a boulder (possibly loose) of weathered meta-sandstone that reveals a dense area (lower right of outcrop) and indistinct bedding with resistant grains (probably quartz) set in a weathered matrix of darker material. Plate C, a photo taken less than a mile from Plate B, shows a small boulder comprising mixed grain sizes, with quartz inclusions (circled in yellow). The heterogeneity of this sample indicates that these sediments weren’t buried deeply enough for recrystallization to occur AND they were close enough to the magma to be injected with fluids .

These samples were observed near the boundary with the Ordovician granite (see inset map). These same sediments look very different further from the granitic intrusion.

Figure 5. Photo of the Cambrian metasediments further from the boundary with the Ordovician granite. The topography here (see inset map of Fig. 4) consists of a series of high, steep mountains comprising these resistant rocks.

Figure 6. Images of outcrops (in-place rock) away from the intrusive contact zone. (A) This photo (two-feet across) shows nearly vertical bedding planes, indicated by a yellow line. Note the uniform surface of this quartzite, which retains original bedding, as shown by dark laminae. (B) This photo shows a contact (yellow line) between the quartzite and a dark rock, which cuts across sediment layers (horizontal in the photo). This stratigraphic enigma may be caused by differential weathering along a fracture, long after deposition and subsequent deformation. (C) This close-up (photo is two-feet across) shows quartz veins dispersed within quartzite.

Let’s quickly return to the inset map of Fig. 4. Now that we know that the sedimentary layers were tilted almost 90 degrees during a mountain-building period, it makes sense that the Cambrian sediments are juxtaposed between younger rocks; they were intruded by granite, while younger sediments were deposited on top of them. These older rocks were heated to a higher temperature by contact metamorphism, and injected with magmatic fluids.

Figure 7. This figure shows a high-resolution image (inset) of the area we discussed in this post. The lower part of the map coincides with the Ordovician granite and the upper part reflects glacial carving of the Cambrian metasediments. Granite weathers much faster than quartzite and is less resistant to physical erosion; thus, the bogs we encountered on our drive are located to the south. This weathering occurred after all of these rocks were tilted during the Paleozoic, and uplifted during the Mesozoic; but before glaciers moved southward (black arrow), forming steep hillsides devoid of soil and trees, during the Pleistocene. That’s a time span of 200 million years.

Summary. Mountain building in Northern Virginia was active between 1000 and 500 Ma. Erosion has removed all geological evidence from the rocks I’ve seen in Virginia, until 200 Ma, when Pangea was torn apart; and North America and Europe were created.

About 550 Ma, sandy sediment was being deposited in shallow seas of Southern Ireland, continuing for at least 100 million years while magma collected in the shallow crust, producing granite intrusions for another hundred-million years. During this dark age in Virginia, the mountains grew then eroded in Ireland and a shallow sea formed, filling with sand, silt, and mud; the shells of marine invertebrates collected in this now-quiet marine environment.

Glaciers never reached Virginia but they covered Ireland throughout the Pleistocene, transforming its landscape.

That’s it in a nutshell …

So far.