Yes, I’m even curious about cobbles brought in to make the bed for the trail I ride my bike on. It would be a lot more interesting if I knew where they got these river-rounded stones, but they’re still interesting!

These rocks are from along Silver Bow Creek in Rocker, just south of Rocker Station on the Greenway Trail, near Butte here in Montana USA. The only thing I can say confidently about them is that Silver Bow Creek did NOT deposit them there – they were brought in.

We can tell that because most (maybe all) of the cobbles are rocks that are not present upstream on Silver Bow Creek or any of its tributaries. Some of the rocks are confidently identifiable as to the formation from which they came, and others are not, but even the black and white granite, which MIGHT have come from the Boulder Batholith, doesn’t really look much like the granite from the Silver Bow Creek watershed around Butte.

I labeled one conglomerate “Post-Kootenai Conglomerate” because the large gray clast right in the middle looks for all the world like a piece of the Gastropod Limestone, a thin unit that’s made up in places almost entirely of gastropod (snail) fossils that probably lived in a large, persistent lake. That rock marks the top, the last-deposited part of the Kootenai formation about 115 million years ago, so for it to be broken into a pebble and then incorporated into another rock (the conglomerate in the photo), this rock has to be later than the Kootenai. There are conglomerates in the post-Kootenai Cretaceous Blackleaf Formation, some of which form boulders along German Gulch and Silver Bow Creek west of here, and which crops out in many places in western Montana.

The main thing missing from this pile of rocks is crystalline Precambrian metamorphic rocks, like gneiss or schist. That makes me think this material was mined from some river gravel in the western part of the state where the Precambrian metamorphics did not erode into the streams, so that leaves out rivers like the Jefferson, Madison, Gallatin, and Yellowstone, all of which should have abundant Precambrian metamorphic rocks in them.

Adjacent to these rocks in Rocker are many big angular broken pieces of Butte granite that probably are part of the ballast holding up the old railroad bed. The cobbles in my photo are most likely part of the restoration, brought in to underpin the Greenway trail in this location.

And yes, I know that technically some of these rocks are not cobbles. The “official” size designation for a cobble is larger than a pebble, smaller than a boulder, or between 64 and 256 millimeters. There are some pebbles and boulders here as well, but I think it’s mostly cobbles (I didn’t do a statistical analysis of measurements of the size distribution).

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Note: This essay is a bit different from what you normally encounter here. It’s just a travelogue from 26 years ago, when my dog Cody and I lived at the Indiana University Geologic Field Station in southwestern Montana. I’m sharing it here from my old website which is likely going away in a year or two. Maybe this will provide some mental cooling on a hot summer day (in the Northern Hemisphere!).

The crescent moon, two days short of the new, was bright with earthshine in the pre-dawn silence of a crisp Thanksgiving morning. From a low of minus 15°F (-26°C) the temperature had climbed with the sun into the single digits when Cody and I set out into Yellowstone Park along the closed and snow-covered West Entrance Road. A raven, black against blue in the sunrise stillness, was the only moving thing. He flew above us along the straight length of the road, surveying our walk. A few cross-country skiers were our only other company — we mostly had the road and the Park to ourselves. The road was snowy but nearly bare in a few places near the entrance, but further in it was completely covered but well-packed by Park maintenance vehicles.

We drove down to West Yellowstone with Indiana U. friends Clyde and Liz from Boise, who had come up to the Field Station for a pre-Thanksgiving visit. We joined their traditional Thanksgiving trek to West Yellowstone where they have gone cross-country skiing for several years. On Thanksgiving Day, Cody and I walked up the road while Clyde and Liz attacked the cross-country ski trails south of town.

About a mile east of the entrance station, the road curves close to the Madison River, and several side roads lead to snow-blanketed and empty parking lots looking along the river. Beyond are vistas of Mt. Holmes and the peaks of the northwest corner of the Park, which rise above 10,000 feet. West Yellowstone and the mostly level entrance road lie at about 6900 feet. Mt. Holmes itself is a Tertiary intrusive, but beyond it are Paleozoic outcrops, including Mississippian carbonates that have been dissolved in the subsurface to produce the calcareous terraces at Mammoth Hot Springs, which is just beyond the hills. These peaks are the southern limit of the Gallatin Range.

The river flowed swiftly and was icy only along snags and logs near its banks. There, every grass and reed that raised its stalk above the waterline slowed the water enough to let an ice-shelf form. But the river put a great breath of moisture into the air which condensed on trees, grass, and weeds. Everything was coated, even the biggest trees. More exposed branches were more densely encrusted, but even pine needles had caught the rime. As befits a bigger stream, the hoarfrost was far more extensive than its counterpart along the South Boulder. And quite possibly the Madison’s water is warmer — we were less than 10 miles downstream from Madison Junction, where the Firehole and Gibbon Rivers come together to form the Madison. Both of those upper tributaries drain major thermal areas of the park.

After our 4-block walk through town to the Park entrance, I had unleashed Cody to romp freely, though I kept him on a fairly close verbal rein. It turned out that this was illegal — when we were almost back to the gate, a ranger out jogging told me to leash him, for his own good: a wolf pack had been spotted within a mile of the entrance just the week before. Neither the first nor last time I’ve violated federal law, but we walked the last half-mile with Cody straining and lunging at the rope, with me wishing I was on a sled that he could pull! But for most of the hike he happily explored the woods within a few hundred feet of the road. When we encountered skiers, he contained his enthusiastic desire to greet them (at my command of “go slow,” which is our equivalent of “heel”) and walked with me.

The state line is about 3 miles in from the Park boundary, and we went nearly two miles beyond that. Most interesting were the side roads to spots that in summer would be congested scenic vistas. For us, they were private entrances to the delicate ice flowers along the riverbanks. Many of the trees were so heavily coated with frost, they looked like flocked Christmas trees, sharply contrasting with the cobalt-blue sky in the background. Among the most fascinating were tall bare trees that were coated with rime right up to the tops.

At almost five miles in, we encountered one of the westernmost areas of burned trees from the fires of 1988. This happens to be one of the areas that was truly burned, killing most of the trees within the fire area. The Park Service has built a boardwalk trail through the margin of the burn zone so that one can compare it to a healthy forest. The boundary between the two is quite sharp. One cross-country skier had skated down the boardwalk before us — and I must admit that Cody did not stay on the wooden path, yet another violation of Federal law.

We saw few rocks on the trek other than boulders along the banks of the river. The Madison here is entering a relatively low, flat area that is probably a basin-range graben related to the same extension and faulting that broke in 1959 to cause the Madison Slide. The substrate below Quaternary sediments is almost certainly Plateau Rhyolite, which crops out in the hills south of the river (and mostly invisible from our traverse). Here, we were beyond the edge of the caldera, but that edge lies near Madison Junction, only about five miles further up the road from where Cody and I turned back.

As always, the return trip seemed a lot faster than the outbound expedition. The walk totaled about 9½ miles and took us about 5 hours. The temperature was about 18°F (-8°C) when we got back to town — to a West Yellowstone that seemed a ghost town compared to the congestion of summer. Within a few weeks, snowmobile season would begin and the streets would be full again — though, perhaps, not for too many more years, given the planned phase-out of snowmobiles in the Park. It was nice for us to have a quiet time both in the park and in town. Cody slept in the car, since the motel had a strict no-pets policy, but he seemed to tolerate that quite well. I think he sees the car as a “secure” area, an extension of the territory of his “pack” — meaning me.

The day was capped by a Thanksgiving dinner that couldn’t be beat in a motel room with Clyde and Liz and friends of theirs from Idaho, Oregon, and Colorado. The next day Cody and I had a leisurely return drive to the Station, catching a view of three moose along the road south of Ennis. A new inch of snow, the first in more than a week, had dusted the northern Tobacco Roots, but the sun was bright and welcoming.

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The white layers in the wall of the Pishkun Canal, north of the Sun River and west of Choteau, Montana, are volcanic ash falls (probably altered to bentonite or some other clay) within the Cretaceous Marias River Shale, which here is otherwise mostly black shale. The incompetent (weak) shale has broken along thrust faults that created folds as well. The thrust faults placed massive strong Paleozoic carbonates (several kilometers to the west (left) of this photo location) over these weak shales, so it’s no surprise that the shale breaks and bends in complex ways.

The Marias River Shale is Late Cretaceous, probably about 75 to 100 million years ago, and the folding and faulting that deformed it is later, broadly 50-70 million years ago.

If you look closely, you’ll see that the sense of offset on the faults appears to be normal, with the left (west) sides down relative to the equivalent layers to the right of each fault. That’s very much unexpected in the Montana Disturbed Belt, which is a zone of compression in which the rocks are mostly pushed to the east, with older rocks up and over their counterparts in reverse faults. These same layers have also been interpreted to indicate east-dipping thrust faults; thrusts are expected, but they’d be expected to dip to the west.

Four of the possibilities (not all mutually exclusive nor exhaustive) here are

1) faults that were originally reverse faults slipped back in the opposite direction after the compression ended, making normal faults dipping west, or

2) there’s been some kind of complicated rotation, possibly as the entire scene visible here rides on a deeper thrust fault, or

3) these faults are conjugate or antithetic faults in a larger fold-fault system whose dominant features we cannot see, or

4) it’s a complex west-vergent (east-dipping) thrust fault within a fault duplex or triangle zone.

(see Bradway, 2007, Stratigraphy and Structural Geometry at the Leading Edge of the Montana Thrust Belt, East of Sun River Canyon, Lewis and Clark and Teton Counties, Montana: U of Montana Master’s thesis; also Sears and others, 2005, Rocky Mountain foothills triangle zone, Sun River, Northwest Montana: 30^th Annual Field Conference, Sun River Canyon and Choteau, Montana: The Tobacco Root Geological Society, Northwest Geology, 34:61, R.I. Gibson and R.C. Thomas, eds.).

This regional map shows the position of the Pishkun Canal photo. Dark green is the Cretaceous shale; the lines with triangular tick marks are thrusts faults that moved from west (left) to east. From Mudge, 1972, Structural Geology of the Sun River Canyon and adjacent areas, northwestern Montana: USGS Prof. Paper 663-B.

The Indiana University geologic field station folks in the photo at top include Alice Nightengale, the late Bob Cassie, and Peter Drzewiecki, which makes this photo the Glacier field trip in 1989.

Pishkun is from a Blackfeet word for “deep blood kettle,” which came to refer to buffalo jumps used by Native American First Peoples. The 12-mile-long (19-km) canal was constructed in 1907-1908 as part of the Sun River Project, a massive effort to bring mountain water to irrigate the plains adjacent to the Rocky Mountain Front. It is also known as the Fort Shaw Canal.

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“Crystal habit” doesn’t necessarily mean the shape and appearance crystals habitually have. It can also describe a rare but distinctive form, or an appearance specific to a particular location.

I’m fascinated by mineral habits because they drive the question “Why?” – why does a mineral grow in one habit here and a very different habit there – sometimes in the same deposit?

The flat lath-like crystals in the top photo are brookite, one of the polymorphs (“many forms”) of titanium dioxide. The other two common ones are rutile and anatase; there are others that are exceptionally rare. Polymorphs are minerals with the same chemistry, in this case TiO₂, but different crystallography.

The crystals in the top photo are very thin – less than half a millimeter for translucent brown crystals as much as 15 mm long. This is quite a distinctive form for brookite at this locality in Pakistan.

The stubby sub-metallic black, almost square crystals below are also brookite. I think a normal person would be surprised to see such diversity in the appearance of the same mineral, but they are both definitely brookite. The examples below of brookite crystals on quartz are also distinctive for their locality, Magnet Cove, Arkansas USA. Magnet Cove has appeared in The Geologic Column several times, including this post on perovskite.

So why are the two examples so different? Crystal habit can be influenced by such things as crystal structure, including preferred growth directions and internal symmetry; unequal or preferential development of various crystal faces; temperature and pressure; space constraints; nutrient supply and crystal growth speed; and impurities. I know of no studies suggesting the origins of the habit in either the Pakistan or Arkansas brookite specimens – if my readers do know of something, please let me know!

Crystal habits include (but are not limited to) words like cubic (and others named for predominant crystal forms), prismatic, fibrous, bladed, tabular, granular, massive, radiating, sheaf-like, stubby, botryoidal, dendritic, arborescent, columnar, acicular (needlelike), equant, micaceous, blocky, stalactitic, globular, reniform, “well-formed,” and twinned, plus distinctive forms named for places where such habits were abundant or first noted, such as Muzo and Tessin.

Sometimes picking a habit name can be pretty subjective, and often enough there’s no sharp boundary between some of them. I’d probably call my crystals from Pakistan “lath-like,” but others might call them bladed. The Magnet Cove crystals might be stubby or blocky, or just something like “well-formed pseudo-cubic.” Sometimes people invent descriptions that are (we hope) self-explanatory, such as “bow-tie crystals” or something else.

It’s even possible for things to get more complicated, as with the example below, which is rutile (tetragonal titanium dioxide) as a pseudomorph (“false form”) replacement of a group of intergrown brookite crystals (orthorhombic titanium dioxide). The rough symmetry might suggest that the original brookite crystals were twins, or maybe it was just several stubby brookite crystals growing together.

Henry James Brooke (1771-1857), brookite’s namesake, was a successful English textile manufacturer and merchant involved in the South American wool trade, but he is remembered as a prolific “amateur” mineralogist who discovered 12 minerals. He is also credited as the first to define the six primary crystal systems.

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Selenium isn’t an element we think about often, although it’s important in several aspects of modern life. And if things had worked out a little differently, it might have been one of the most important elements we use.

Selenium was discovered in 1817 by chemists Jöns Jacob Berzelius and Johan Gottlieb Gahn who also owned a plant to produce sulfuric acid at Gripsholm, Sweden. They thought a red precipitate, a byproduct of processing pyrite for the sulfuric acid, was a tellurium compound, but eventually discovered a new element they named selenium, for Selene, Greek goddess of the moon, by analogy with tellurium, named from Latin tellus, Earth.

But it wasn’t until 1873 that Willoughby Smith (1828-1891), working on electrical properties of materials for the company that made undersea cables for telegraphy, discovered selenium’s sensitivity to light. I recounted his story in What Things Are Made Of (p. 177-178), but once his accidental discovery that selenium’s electrical properties depended strongly on light was confirmed, all sorts of technologies opened up.

Astronomers Joel Stebbins and Jacob Kunz at the University of Illinois used selenium compounds to image the faint light of distant stars in the early 1920s. And in 1920, A.O. Rankine published a note in Nature (Feb. 5, 1920) titled Telephony by Light. He showed that a “selenium valve” could be controlled by light, and fluctuations in light produced by the vibrations of a diaphragm by sound could be transmitted in a light beam to a selenium receiver and converted back into sound. He recognized that Alexander Graham Bell had invented a photophone in 1880, successfully transmitting sound using a light beam over 200 yards.

Rankine’s apparatus used six-inch lenses and he reported that “the articulation of the speech heard [was] extraordinarily perfect” over a distance of 8 miles. He felt that the range was limited only by sight distance, ultimately the curvature of the earth – and of course this was long before lasers or satellites. Nonetheless, obviously photophones didn’t catch on, at least not those using selenium, and Bell’s telephone based on the electrical properties of a magnet moving in proportion to sound waves came to dominate the industry.

Historic and traditional uses for selenium included decolorizing glass, an application that still commands about 20% of selenium consumption. The greatest use (40%) of selenium is in processing manganese and some other metallurgical work, making electrolytic production less consumptive of electric power.

Historically, photovoltaic selenium compounds were used in camera light meters and standard photocopiers. In 1995 in the US, about a third of the selenium went to photocopiers, but by 2000 it was near zero as more efficient organic-silica compounds were developed. The same is probably true of electric-eye doors that formerly used selenium-based materials. Copper-indium-gallium-diselenide (CIGS) is still used in solar cells. With other electronic uses this adds up to about 10% of selenium consumption.

Selenium is a vital component in human and animal health, although it can concentrate deleteriously in the environment. Its use in fertilizer and livestock feed amounts to about 20% of the total globally. It can also be found in anti-dandruff shampoos.

Selenium is rarely concentrated in specific minerals, although there are some. My micromount examples in the top photo are definitely not in the category of “spectacular” in appearance.

Virtually all selenium (and tellurium) is a by-product of electrolytic refining of copper, so its production is strongly dependent on the copper industry worldwide. In the 1990s refineries in Japan and Belgium yielded more than 50% and 13% respectively of world selenium, but in 2025 the unsurprising leader is China at 53%, reflecting both their copper refinery production (48% of the world) and position as the world’s leading copper consumer. But Japan is still in second place, at 17%. Other leaders include Russia (8%), Belgium (5%), and Canada (3%). Production statistics from two US refineries (in Texas and Utah) are withheld to preserve proprietary data, but the US is probably close to 90% (or more) dependent on selenium imports, from South Korea, Philippines, Mexico, Chile, China, Poland, and elsewhere.

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Extinct ammonites comprised a subclass of cephalopods (“head foot”), part of the phylum of mollusks. The only surviving members of the cephalopods, octopuses, squids, cuttlefish, and nautiluses, are actually not that closely related to ammonites and are in a separate subclass.

There are many subdivisions among ammonites, but probably the most familiar are coiled and straight shells. Baculites, a straight-shelled genus, was abundant in Montana (and worldwide) during the Cretaceous period. The two in my photos above come from the Bearpaw Shale of eastern Montana, and date to the Campanian Stage, which ranges from about 83 to 72 million years ago.

The coiled ammonite at top is probably Placenticeras and this one is probably from the Pierre Shale of South Dakota (I got it in Rapid City in 1970) but it is identical to examples from Montana. The Bearpaw Shale is the equivalent of the upper part of the Pierre Shale. Both shales were deposited in the Cretaceous Western Interior Seaway of North America.

The Western Interior Seaway was a shallow sea that developed in response to the rising highlands in western Montana, Idaho, and Utah, and extending to the north and south. The crust east of the collision zone that created the highlands was depressed beneath the rocks that were tectonically thrust upon it and was further depressed by the pile of sediment shed off that highland into the adjacent lowland (the depression is called a foreland basin).

The seaway began to develop about 100 million years ago and ultimately extended the length of the tectonic orogenic (mountain-building) belt from the Arctic to the Gulf of Mexico, and at times it was as wide as from western Montana to Minnesota, and New Mexico to Arkansas.

The name “seaway” makes it sound like it was just an ocean, but the system included a wide range of environments, which you can see on the map above from the USGS. From the eroding mountains and volcanoes in the west, streams flowed eastward onto a wide alluvial plain, and the associated rocks record stream channels, mud flats, lakes, and alluvial fans there. Further east the low coastal area included (at diverse times) sandy beaches, lakes and coal swamps, and deltas of the rivers. The true marine environment, where the ammonites lived, was even further east. Much of the sea received fine-grained mud as sediment, which ultimately solidified into the Bearpaw and Pierre Shales. When ammonites and other swimming creatures in the sea died, they settled onto the muddy sea floor where they were preserved.

The Western Interior Seaway lasted around 40 million years. When it had lasted about 30 million years, around 70 million years ago or so, ongoing tectonic activity began to break the crust beneath the seaway’s basin. Huge uplifts, which persist today in the Uinta, Wind River, Big Horn, Beartooth and other mountain ranges, and in the Black Hills, began to segment the basin and change the depositional patterns. Ultimately the floor of the basin was above sea level, roughly 60 million years ago, and marine deposition ceased.

Ammonites had gone extinct before the demise of the Western Interior Seaway. You commonly hear that they expired in the end-Cretaceous extinction, the same event that marked the end of the non-avian dinosaurs, but there is some controversial evidence that at least one Baculites species may have survived perhaps 200,000 years into the Paleocene, after the extinction event (Machalski and Heinberg, 2005, Evidence for ammonite survival into the Danian (Paleogene) from the Cerithium Limestone at Stevns Klint, Denmark: Bulletin of the Geological Society of Denmark, Vol. 52, pp. 97–111).

Ammonites are typified by their intricate and beautiful suture patterns. The convoluted patterns called sutures are the intersection between the internal septa, layers bounding the chambers within the shell, and the layer just below the outer surface. My coiled specimen does not show the sutures because it still has much of the original outer shell material (nacre), which is iridescent in places. The Baculites specimens show the sutures well.

We don’t fully understand either how the suture patterns developed, nor exactly what their purposes were in the living animals. Except for the initial chamber, where the animal lived, the chambers in the shell were empty, and it is usual to infer that the animal changed the water-air ratios in those chambers for propulsion and buoyancy. But how the exceptional convolutions of the septa dividing the chambers developed is still being researched.

Inoue and Kondo (2016, Suture pattern formation in ammonites and the unknown rear mantle structure: Nature, Scientific Reports 6, no. 33689) used the almost fractal projections that develop on modern sea slugs, unshelled mollusks, as analogies for the growth of ammonite bodies that would control the development of the septa and sutures. However they formed, the suture patterns are so distinctive that ammonites can be used to establish high-precision relative dating of stratigraphic layers, especially during the Mesozoic era.

“Baculites” is from Latin and Greek words meaning “walking stick rock.” Some varieties reached two meters in length so could indeed have served as walking sticks although they tend to break apart into smaller segments along the septa. Several Baculites species from Montana’s Bearpaw Shale were described by William A. Cobban, a USGS geologist and paleontologist descended from the same family whose name is commemorated in Cobban Street in Butte, Montana (Cobban, 1962, New Baculites from the Bearpaw shale and equivalent rocks of the Western Interior: J. Paleo. 36:1, p. 126).

“Ammonite” derives from Pliny the Elder (who died in the 79 AD eruption of Vesuvius) who called the fossils ammonis cornua (”horns of Ammon”) because the Egyptian god Ammon or Amun was typically depicted wearing rams’ horns that looked similar to coiled ammonites. The common suffix in ammonite names, -ceras, as in the Placenticeras specimen in my photo, also means “horn.”

We’ve looked at aspects of the Western Interior Seaway of North America previously, here and here.

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John Day (c. 1770-1820) was a trapper and hunter in what is now the Pacific Northwest. He is recalled in the names of rivers, towns, and the John Day Fossil Beds National Monument.

The rocks in this region of central Oregon USA range from Eocene to Miocene in age, about 45 to 5 million years ago. Throughout much of this time volcanic activity, ranging from lava flows and lahars (volcanic ash-mud flows, from the Javanese word for these features and events) to extensive ash falls erupted into a lush, sub-tropical setting.

The volcanics resulted in excellent preservation of the plants and animals there over a span of tens of millions of years, ranging from diverse plants to large mammals such as brontotheres (relatives of rhinoceroses) and extinct relatives of hyenas. Brontops in the image above is now named Megacerops, and is a kind of brontothere. This drawing is from a c. 1905 trading card published by a German chocolate company. More information about that in this previous post.

The dramatic color variations in the rocks, both the volcanics and the sediments derived from them, reflect both variable composition and variable climatic conditions, resulting in diverse oxidation of the iron in the rocks. Iron is by far the strongest and most common coloring agent in rocks, and most of the colors here relate to iron content.

Bright red indicates oxidized iron (rust) and is interpreted to have developed during very wet (even swampy), warm conditions. A change to yellow and tan represents a somewhat drier environment, but still probably considerably wetter than this area is today. Occasional thin black zones are probably relatively rich in manganese oxides, possibly related to fixing in the soils by plants. Most of the layers of all colors are essentially paleosols, old soils that developed between volcanic events, but specific ash falls and lava flows are certainly present as well.

The earliest volcanism (the Eocene Clarno Formation, 53-43 million years ago) was not related to the Cascades which did not begin until about 37 million years ago with major activity from around 30 to 17 million years ago. Today’s Cascade Range reflects a rejuvenation of volcanism mostly over the past 9 million years.

The volcanism that produced the Clarno Formation appears to have related to pre-Cascades subduction (Bestland and others, 2002, Geologic Framework of the Clarno Unit, John Day Fossil Beds National Monument, Central Oregon: Oregon Dept. of Geology and Mineral Resources Open-File Report O-02-03). The Clarno volcanoes may have developed after an exotic terrane called the Siletzia Block collided with the margin of North America, jamming the subduction zone, steepening its angle of descent, and breaking the overlying crust to allow the Clarno volcanism (Miller, 2014, Roadside Geology of Oregon: Mountain Press Publishing, now handled by the Geological Society of America).

My photos are from a visit in 2019. No color enhancement.

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If you’ve been around The Geologic Column a while, you know I sometimes acquire cheap specimens without knowing where they are from, for the enjoyment of trying to figure it out. This time, it’s a known location and an uncertain mineral.

The specimen was collected by Duane Johnson probably in the late 1980s at the Black Pine Mine near Philipsburg, Montana (you’ve seen me ramble on about Black Pine repeatedly, too, and I imagine there will be more). Duane held onto it and many others until this year, when he sold it at the Butte Mineral Show. I spent the princely sum of $2 for the piece, labeled pseudomalachite, copper phosphate, Cu₅(PO₄)₂(OH)₄, a common mineral at Black Pine.

It’s certainly not pseudomalachite, which at Black Pine typically forms distinctive deep green spheres and botryoidal masses, or when rarely crystals, usually intergrown lath-like blades.

Obviously X-ray or Raman analysis could probably provide the identity of these crystals, but for me it’s much more fun to go as far as I can on my own.

Crystallographically and chemically, the primary suspects are the minerals in the list at the top. Even though the last three are not reported from Black Pine on MinDat, their chemistry means it would not be even slightly surprising for them to show up there. But I’ll focus on the first three in the list for now.

Philipsbornite is not a combination of philipsburgite, type locality at the Black Pine Mine, and bornite, common at Butte. Rather it was named in 1982 for Helmuth Richard Hermann Adolf Friedrich von Philipsborn (1892-1983), professor of mineralogy at the University of Bonn. Its type locality at Dundas, Tasmania, is probably the most famous locality for crocoite in the world.

But it was this specimen from South Africa that struck me as quite similar to my specimen. Although the crystals in my specimen may look like little octahedrons, it doesn’t take much looking to see that they are not – they are pseudo-octahedral, a fairly common habit for philipsbornite, also seen in this example from Germany. The crystals are (I think) stubby, equant trigonal trapezohedrons. There are no photos of philipsbornite on MinDat from Black Pine but it is listed there. It has a huge color range depending on impurities (mostly iron and copper), from colorless and white to yellows and greens to bluish.

A specimen of duftite that’s similar to mine can be seen here. Duftite often makes equant orthorhombic dipyramids and similar appearing orthorhombic disphenoids. The few examples of Black Pine duftite on MinDat are gemmy bright green crystals, and others I have seen are like those only darker.

Olivenite is seen in only a few photos from Black Pine on MinDat. Crystals that I have from Black Pine show typical flattened laths, doubly terminated with wedge-shaped faces.

Philipsbornite and duftite both have a hardness around 4.5, and olivenite is 3.0. It’s challenging to get a good hardness on my tiny crystals, but such as I can determine, they do seem to be on the harder side, arguing against olivenite.

So my working identification for these little crystals is philipsbornite – although admittedly I want it to be the more obscure of the choices so that may well prejudice me. I’ll update this page if and when I get analytical confirmation.

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A terrane is a package of rocks that share a common geologic history and has typically moved great distance from its origin and has been accreted (added tectonically, “sutured”) to another, often larger, tectonic plate. Terranes can be highly diverse, containing fragments of continental crust, volcanic island arcs, oceanic igneous provinces, oceanic crust, standard sedimentary rocks, and more. They are usually bounded by faults or fault zones, but they are typically much larger than simple fault blocks, usually on the order of hundreds of kilometers or more in length.

A terrane could be something as complex as Japan, which probably includes some Precambrian continental (or continental margin) rocks, oceanic crust, lots of volcanic rocks related to early and present-day subduction, and sedimentary rocks derived from all of those. Or it could be relatively simple, like the long string of mostly basalt that forms the Hawaiian Island chain.

The folds and little faults in the photo at top are in Silurian carbonates (deposited around 420 million years ago) within the Alexander Terrane in Glacier Bay, Alaska (Blodgett and others, 2013, Sapelnikoviella santuccii, a new gypidulinid brachiopod genus and species from the upper Silurian of Glacier Bay National Park & Preserve, Southeast Alaska: Memoirs of the Association of Australasian Palaeontologists 44:65-72).

The Alexander Terrane is usually considered to be related to the Wrangellia (Vancouver Island and elsewhere) and Peninsular Terranes (southwestern Alaska). While the three terranes are distinct, they seem to have been in relatively close proximity to each other when they were attached to North America probably sometime during the Jurassic Period, about 150 to 165 million years ago, or as recently as early Cretaceous, 85 to 110 million years ago. The three are sometimes called the Wrangellia or Insular Composite Terrane, which lies between older accreted terranes further inland in Alaska, Yukon, and British Columbia, and younger terranes closer to the Pacific Ocean.

The tilting and small folds and faults in the photo at top could have resulted when the Alexander, Wrangellia, and Peninsular Terranes interacted with each other on approach, or more likely when they all collided with the older margin of southern and central Alaska. The cross section above shows the relative position of the Alexander-Wrangellia-Peninsular Terrane (APW) between older terranes to the north (right) and younger to the south.

The three photos immediately above show multiple dark dikes cutting across the layers of Silurian carbonates. I found no reports about these dikes, which by their dark color suggest basaltic composition, but because they are relatively undeformed, I’d guess that they are probably related to relatively recent igneous activity associated with modern subduction and faulting - the rising red blobs of magma shown in the volcanic arc on the cross section. At least that’s what I speculated to the group on the Smithsonian Journey in 2005 when I was their house geologist and took these photos.

The location of my photos is called Jánwu Aaní in the Huna Łingít language, meaning “Mountain Goat’s Land,” although when I was there we didn’t see any mountain goats. It is also called “Gloomy Knob.” On the geologic map above, the Silurian rocks are dark blue-gray, labeled DSsm.

Where did the Alexander Terrane (and the others) come from? Evidence is quite contradictory, but one recent study (Beranek and others, 2013, Baltican crustal provenance for Cambrian-Ordovician sandstones of the Alexander terrane, North American Cordillera: evidence from detrital zircon U-Pb geochronology and Hf isotope geochemistry: Journal of the Geological Society, 170 (1), 7-18) suggests that it originated at least 500 million years ago in a complex island arc terrane off Baltica (the European Craton), north of Norway in modern coordinates. It reached Alaska in a sliding, transform manner somewhat like the push of the Caribbean Oceanic Plate and the associated continental fragments (like Cuba) and volcanic arcs (the Antilles) between North and South America. That movement is the origin of the recent devastating earthquakes in northern Venezuela. But for the Alexander Terrane, the two continental masses it was squeezed between were Europe and Siberia, before they came together. The details of Alexander’s motions after about 425 million years ago, into what is now the Pacific realm, remain enigmatic, as far as I can tell.

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These little greenish worms are clinochlore, or at least somewhere in the chamosite-clinochlore series, more generally known as chlorite. Clinochlore is a magnesium aluminosilicate that’s pretty typical in low-grade metamorphic rocks. The worm-like elongate things are essentially long books of thin sheets parallel to each other, aligned to make the “worms.” The color variations from light to dark green may indicate differences in iron content.

Clinochlore is Mg₅Al(AlSi₃O₁₀)(OH)₈ and chamosite is Fe²⁺₅Al(AlSi₃O₁₀)(OH)₈. Without analysis, most geologists would probably just call this chlorite. Clinochlore was named in 1851 from Greek κλινειν klinein, “to incline,” referring to inclined optic axes visible in thin section, plus Greek χλωρο chloros, for “green,” its typical color. Chamosite’s name goes back to 1820, for the type locality Haut de Cry, near Chamoson, Valais, Switzerland.

The white substrate is the mineral meionite, Ca₄Al₆Si₆O₂₄CO₃, the calcium end member of the scapolite group of aluminum silicates.

The rock is from an outcrop along the Pipestone-Delmoe Road east of Homestake Pass, east of Butte, Montana, USA, where the Butte and Rader Creek Plutons of the Boulder Batholith trap between them a zone of formerly sedimentary and/or volcanic rocks that were metamorphosed and metasomatized (which means there was chemical exchange) by their contact with the granites of the batholith. The zone, informally called the Screen, contains some interesting complex mineralogy because the chemicals in the rock were remobilized by the heat from the molten rock of the batholith about 76 million years ago.

We’ve visited the Screen in this previous post, and there’s some more information in my paper in Scarberry, K.C., and Barth, S., editors, 2018, Proceedings of the Montana Mining and Mineral Symposium 2017: Montana Bureau of Mines and Geology Open-File Report 699.

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The broadest definition of fossils is probably “evidence of past life.” That can be taken to mean trace fossils (things like trilobite tracks) and unusual chemical signatures as well as actual body fossils. Whatever molar-tooth structures are, they are NOT related in any way other than (maybe) appearance to canines, incisors, molars, or teeth of any sort. And it’s uncertain whether they may be indicators of past life.

What it is: Molar-tooth structure is a pattern in Precambrian rocks, mostly carbonates, of cracks and open spaces usually filled with microcrystalline calcite (called microspar; spar refers to crystals with sharp faces or cleavages, from Old English spærstān, crystal projections). In practice, the crack filling and matrix are clearly and sharply delineated, with gray, sparkly microspar in the cracks and lighter color, granular matrix. The microspar may occupy highly convoluted sheets and ribbons or irregular to rounded blobs (Kriscautzky and others, 2022, Molar-Tooth Structure as a Window into the Deposition and Diagenesis of Precambrian Carbonate: Annual Review Earth and Planetary Sciences. 50:205-230).

Molar-tooth structure is known in Precambrian sedimentary rocks from around the world since its first discovery on the Belt Supergroup rocks of southeastern British Columbia and adjacent Montana in 1885. Its origin has been debated since that time. The structure is restricted to Precambrian rocks, with no modern analog, so understanding it may lead us to recognizing processes in the early Earth that at present are obscure.

Kriscautzky and others (2022) offer this summary of potential origins:

1. Biological origin – actual fossils, representing remnants of algal growths or some atypical aspect of stromatolites (algal mats).

2. Relicts of evaporites, such as anhydrite or gypsum, that formed in shallow unconsolidated sediment, were dissolved, and were later replaced by calcite and deformed.

3. Shrinkage features in solidifying carbonate mud – essentially deformed mudcracks or other dehydration features developed as the sediment consolidated into rock.

4. Cracking related to stress with later fill by calcite. Partially consolidated mud may move and deform as soft sediment when stress is applied. The stress could come from liquefaction under gravity (wave or tidal loading) or as a result of seismic activity, or simple enhanced fluid flow as parts of the sediment package solidify at different rates.

Kriscautzky and others (2022) indicated that all of the above ideas are inadequate to explain all aspects of molar-tooth structure. The favored explanation today is some kind of gas expansion and migration. The gas, whose nature is still debated actively, but which was most likely carbon dioxide, hydrogen sulfide, or methane, would have been generated by microbial decomposition of organic matter. The primary life form during the time of Belt Supergroup deposition about 1,400 million years ago was algae and cyanobacteria – but tiny though they were, they were enough to oxygenate the earth’s atmosphere over several hundred million years close to a billion years before the Belt rocks were deposited.

Gas migrating through soft sediment, partially consolidated, could produce the void spaces as well as the linear tracks and ribbons observed in molar-tooth structure. Gas escape could have ranged from gradual to explosive, which together with complex variations in the sediment (strength, grain size, porosity, and ongoing soft-sediment deformation) can explain the exceptionally diverse geometry of molar-tooth structure. At least the consensus appears to be that that is the best explanation to date.

If the gas escape interpretation of molar-tooth structures is correct, they do serve as “evidence of past life,” albeit microbial.

My specimen is from the Helena Formation near Helena, Montana, which is the equivalent of the Siyeh Formation in Glacier National Park where molar-tooth structure is displayed spectacularly in some of the rocks supporting the wall around the parking lot at Logan Pass (photo below).

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In technical geologic and tectonic terms, Inner Mongolia is a mess. It’s been collided, extended, intruded, and extruded so much, with an overprint from the relatively recent collision with India thousands of kilometers away, that it’s an area at the forefront of detailed tectonic understanding. Chinese scholars are driven by the presence of many diverse economic mineral deposits.

There’s more than one Huanggang in China. The name means “yellow hill,” of which there are probably plenty in China, with iron oxides coloring the surfaces. My specimen at the top of this post is from the one in Inner Mongolia, northeast China, home to an important iron-tin deposit.

Huanggang is famous with collectors for its fluorite, ilvaite, andradite, and arsenopyrite. The hematite in my specimen would probably be counted among the ore. The deposit is a skarn, affected both by metamorphism (heat and pressure) and by metasomatism (exchange of chemicals between country rock and intrusive rock). The age of the granitic intrusive and the mineralized skarn is about 132 million years, Early Cretaceous time (Xue and others, 2025, Ore Genesis of the Huanggang Iron-Tin-Polymetallic Deposit, Inner Mongolia: Constraints from Fluid Inclusions, H–O–C Isotopes, and U-Pb Dating of Garnet and Zircon: Minerals, 15(5), 518. https://doi.org/10.3390/min15050518). The country rock is mostly Permian sedimentary rocks, much of it now metamorphosed to marble.

The igneous and tectonic activity that produced the Huanggang and many other mineral deposits in Inner Mongolia was ultimately related to a complex collision called the Central Asian Orogenic Belt. It represents an amalgamation zone more than 1,000 km wide between the dominant Siberian and North China Cratons. The zone includes many smaller terranes (called “microcontinental blocks” on the map above) separated from each other by regional fault systems that undoubtedly contributed to the fluid movements that produced the mineral deposits. The Central Asian Orogenic Belt has a very long history of deformation extending back at least to Carboniferous-Permian time (300 million years ago) with later major activity in the Early Cretaceous (132 million years), giving rise to Huanggang and dozens of other mineral deposits.

The hematite (iron oxide) in my specimen shows hundreds of stacked plates of hematite which together make rounded aggregates that are steely blue-black in color. The associated quartz forms a sheaf-like group with a very pale purplish tint in parts.

China is the #4 leading miner of iron ore in the world, after Australia and Brazil and just slightly behind #3 India (the US ranks tenth in iron ore mining). Much of the iron production from Australia, Brazil, Canada, the US, and elsewhere is from banded iron formation, mostly just iron oxide and quartz. China’s iron deposits are polymetallic, enriched in a wide diversity of metals including copper, lead, zinc, silver, tungsten, tin, rare-earth elements, and more, making them much more than simple iron deposits.

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Rattlesnake Butte, South Dakota, is a well-known locality for calcite crystals so heavily included with sand that they are called sand crystals. The locality today is a National Natural Landmark, managed by the Oglala Sioux Parks and Recreation Authority on the Pine Ridge Indian Reservation, and collecting is prohibited without a permit.

The sandstone is Miocene in age and is compositionally very immature, meaning it contains grains of many minerals and rocks besides simple quartz. Wanless (1922, Notes on sand calcite from South Dakota: Amer. Mineralogist, 7:83-86) reports “quartz, orthoclase, microcline, albite, muscovite, biotite, hornblende, augite, green, black and colorless prisms of tourmaline, zircon, garnet, staurolite, magnetite,” and rock fragments of volcanic glass and mica schist. All that indicates the sand was derived largely from the pegmatite-rich Precambrian rocks of the Black Hills, which lie about 80 miles to the west.

My example here has definitely suffered some post-crystallization erosion or dissolution. You can find examples in which the crystallography of the calcite remains sharp and distinctive.

The sandstone is part of the Arikaree Group, the youngest rocks in Badlands National Park which lies north of Rattlesnake Butte. The sand was probably wind-blown.

The timing of the formation of the calcite crystals is uncertain. Wanless (1922) suggested they formed in the subsurface, from percolating hot-spring water or groundwater, but either way, the water must have carried significant dissolved calcium and carbonate to precipitate the calcite within what must have been very poorly consolidated or even loose sand. But Cirone and Law (2005, Microstructure of calcite sand crystal and implication on its crystallization process: Geological Society of America Abstracts with programs, v. 37, no. 1, p. 59) felt that sand grains floating in the calcite could not be explained by simple subsurface crystallization.

Calcite composes about 35% to 40% of the crystals, with the remainder the sand. The crystals are usually scalenohedrons with some modifications and often rounding from later weathering.

See also Huizing, T. E., & Richards, R. P., 2021, Connoisseur’s Choice: Sand Calcite, Rattlesnake Butte, South Dakota: Rocks & Minerals, 96(5), 414-431, and this post at CSMS Geology.

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Beaverhead Rock along the Beaverhead River in southwestern Montana is famous as the feature recognized by Sacagawea, Lewis and Clark’s Shoshone guide, on August 8, 1805, placing the Corps of Discovery somewhere near the Continental Divide.

The rock of Beaverhead Rock is the Mississippian-age (Early Carboniferous) Madison Limestone, a thick package laid down in mostly shallow, warm water about 330 million years ago. Montana’s Madison Limestone and its equivalents record a seaway that extended from Arizona to Alberta and east to Indiana and beyond. See the map and photos in this previous post for more about the shallow Carboniferous sea in North America.

Apart from its distinctive shape, the outcrop is also prominent because it is isolated, one of few large exposures of rock anywhere near the river over a distance of many miles. The isolation is a consequence of complex tectonics.

The northeast-southwest trending Beaverhead River valley is in a graben, a down-faulted zone resulting from extension, pulling apart, of the entire region. This dates to around 35 to 15 million years ago and to some extent is going on today. The lightest colors on the map above, labeled Qa, represent the alluvium and recent deposits from the modern river system.

The valley is narrowest at Beaverhead Rock. I think that’s because of the northwest-trending fault zone there, highlighted on the map and with a dashed probable extension to the Ruby Range, where numerous northwest-trending faults are mapped. This northwest-southeast structural grain dominates the Tobacco Roots, the Ruby Range, and parts of the Highland Mountains, and probably derives from a very old (Paleoproterozoic, at least 1,700 million years ago) fault system that has been repeatedly reactivated over time up to the present.

The part of the Madison Limestone that’s exposed on Beaverhead Rock is the upper part, the Mission Canyon Limestone, which contains broken (brecciated), collapsed rock in paleokarst. Karst topography develops when limestone dissolves to produce caves and other features; paleokarst is the same thing, but in this case it developed perhaps 300 million years ago, soon after the Mission Canyon Formation was deposited. The rocks must have been uplifted and exposed at the surface where ground water created caves that later collapsed to form the broken rock and cavities we often see at the top of the Mission Canyon Formation.

Cavities now containing large calcite crystals probably developed during a second phase of dissolution and re-deposition that dates to Oligocene-Miocene time when southwest Montana was much wetter than it is today, but the cracks often follow the old collapse structures in the paleokarst. In places, not-very-well-formed calcite crystals reach 6-7 cm like those in the photo above. It’s a State Park so collecting is not allowed.

The calcite crystals in the photo above, left of center, are about 5 cm long and occupy an S-shaped opening in the rock.

The Beaverhead River seen from Beaverhead Rock shows beautiful meanders. This part of the stream system is a relatively low-gradient river, allowing the lazy meanders to develop.

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Galena, lead sulfide, is the primary ore of lead. Native lead exists and looks a lot like galena, but it is softer, malleable, and lacks galena’s distinctive cubic cleavage.

Even though lead ranks #36 in abundance of the elements in Earth’s crust, lower than zirconium or gallium or rare-earth elements like lanthanum and niobium, lead and sulfur are reactive enough that they often form galena, which in turn makes common metallic mineral deposits. Lead was smelted from galena in what is now Turkey about 8,500 years ago – probably not to get lead, but to isolate the silver that was typically associated with it. By Roman times, it was cheap and readily available and was in common use for water pipes. Latin plumbum for lead gives us the word plumbing and lead’s chemical symbol, Pb.

Galena is usually one of the minerals students learn in their first geology classes. It’s heavy, and the lead-gray to silvery-gray (on fresh surfaces) color and sharp cubic cleavage make it easy to identify.

My two specimens here are from the Tri-State District of the US, around the common corner of Kansas (near the town of Galena), Missouri, and Oklahoma. This area was historically one of the greatest producers of lead and zinc, extending from the 1850s to the late 1960s-early 1970s.

The deposit is a Mississippi-valley type deposit in which the ores have migrated in hydrothermal solutions, typically brines, that often originate in deep basins heated by underlying magmatic sources. The mineralized fluids migrate through porous rocks, or along conduits such as fault zones, until they reach conditions of temperature, pressure, and host rock that allow minerals to precipitate. This often happens on the flanks of a regional uplift adjacent to the brine-bearing basin. For the Tri-State District, that regional uplift is probably the Ozarks in Missouri.

Today, the Tri-State District is being reclaimed as multiple Superfund sites, but Missouri is still the leading producer of lead in the US, from the Viburnum Trend on the other flank of the Ozarks in southeastern Missouri.

Together with Alaskan production (where lead is a byproduct of zinc mining) and Idaho (a byproduct of silver mining), the Missouri lead contributes to make the US the 4^th leading lead miner in the world, approximately in a 5-way tie with Peru, Russia, India, and Mexico, each producing about 4% to 6% of world lead. As usual, China leads with 42% and Australia is #2 with 11%.

Almost all US lead mine production is exported, because the US has not had an operating lead refinery since 2013. Export destinations for the ore and concentrates include China, South Korea, Canada, Switzerland, Mexico, Japan, and Australia. Their refined product comes back to the US as part of a 33% reliance on imports, mostly from Canada, South Korea, Mexico, and Australia. The other 67% of US lead consumption comes from recycling.

China and Australia control most of the world’s lead reserves. On the map above, differences in Australia’s reserve number reflect two estimates: USGS at 36% of world reserves, and Australia’s more conservative Joint Ore Reserves Committee at 11%, based on a higher grade than the USGS number.

The primary use of lead is in lead-acid batteries for vehicles, on the order of 88% of total consumption. Recycling those batteries is a big part of the lead industry; US recycling consists mostly of exporting more than 30 million lead-acid batteries per year to the countries listed above. This makes for a complex global movement of lead ores and the metal: Mined and recycled material are exported from the US and other producers, refined in other countries along with their own mine production, then imported to the larger consumers, the US and China, to make more lead-acid batteries.

Other small uses for lead include ammunition, solders, and in glass and ceramics. One historically large end use, lead in gasoline, has been eliminated because of the negative health and environmental effects it caused. Algeria was the last country to ban leaded gasoline, in 2021. Another historical use that has diminished is lead in glass for cathode-ray tubes (as in old televisions) to block x-rays.

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In the fall of 1984, I was unemployed, between quitting Gulf Oil in May and starting with Everest Geotech in November. It was a perfectly good time to take in the Montana Geological Society conference in northwest Montana.

Arco had just finished drilling a very wild wildcat well in the hinterlands west of Kalispell, Montana, to test a radical tectonic concept: the idea was that the sedimentary rocks that are east of the Rocky Mountain Front, and beneath the Lewis Overthrust in Glacier National Park, MIGHT actually be present at great depth beneath thrust faults that had pushed ancient Precambrian rocks over them, much further to the west.

The well took nearly a year to drill, in 1983-84, and while I don’t think the cost is public, rumor at the time suggested it cost a million dollars (or two) or more – huge in 1984. The curved seismic reflectors there MIGHT have represented a huge oil or gas trap, which if present would have justified the cost.

It wasn’t present. They drilled almost 18,000 feet of Precambrian Prichard Formation, the age and type of rock that is extremely unlikely to hold oil or gas fields. The seismic reflectors turned out to be about 3,500 feet of igneous sills. The well effectively ended what was for a time almost an “oil rush” in northwest Montana, making plenty of people happy and some explorationists sad. But it did add to our understanding of the tectonic fabric of the region.

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What is a classic mineral or mineral locality? It’s subjective, of course, but Thompson and others (2025, Ikons of the Mineral World: Nature’s Finest Art: The Mineralogical Record, Inc.) say it refers to “great minerals and to the localities that produced them, usually in the 19^th or early 20^th centuries.” And that the localities are usually extinct, and therefore “True mineral classics are extremely collectible and are also nearly impossible to obtain.” They explicitly include purple apatite from the Pulsifer Quarry in Maine as a classic mineral from a classic locality.

Full disclaimer up front: I’m not absolutely certain that the purple apatite crystals I show here are from Pulsifer. All I know for sure is that they are from Maine, but I think it’s quite likely they are from Pulsifer.

In 1901, Pitt Pulsifer (1850-1934) discovered the first gemmy purple apatite crystals at his workings, where gem tourmaline had previously been found. The crystals were remarkable enough to get the attention of Harvard mineralogists John Wolff and Charles Palache (1902, Apatite from Minot, Maine: Proceedings of the American Academy of Arts and Sciences, Vol. 37, No. 18, p. 517-528). Harvard’s Mineralogical Museum acquired “the whole” of Pulsifer’s initial find.

Of Pulsifer’s material, Wolff and Palache wrote

“This find is noteworthy for the unusually rich purple color of the crystals, and the purity, crystalline perfection, and abundance of the material, which comprises about two thousand loose crystals or fragments of crystals with a total weight of over a kilogramme, and about a dozen large groups of crystals on the matrix. Of the loose crystals about three hundred show at least one perfect termination, five hundred are slightly less perfect, and the rest imperfect or fragmentary.”

The apatite crystals grew in a pegmatite cavity in granite, with quartz, orthoclase, and lepidolite. The prismatic fluorapatite crystals averaged about 10 mm in length. Analysis showed traces of iron, aluminum, magnesium, and manganese, but Wolff and Palache (1902) did not speculate on the origin of the purple color. Today, you’ll see statements that the color is due to manganese.

Pulsifer continued to mine the quarry until 1917. It was dormant until 1939 when Harvard leased it to explore for crystals, finding tourmaline but no apatite. Lessors mined the property in 1964-1967 finding outstanding purple apatite crystals (Thompson and others, 2025).

My third specimen (above) is also from Maine but also from an unspecified precise locality. This stubby hexagonal prism is a little unusual for Pulsifer, where the crystals seem to be somewhat more typically elongate prisms, but it could still be from there. But there are at least 491 photos of apatite from Maine on MinDat, so there are other alternatives. For example, the Emmons Quarry, Uncle Tom Mountain, Oxford County, Maine, has yielded many quite similar to this hexagonal specimen.

Here’s the provenance of my specimens. I got them when my mineralogy professor at Indiana University, Carl Beck, died in 1971, when I also continued his side business analyzing the mineralogy of kidney stones. It’s possible that he obtained the apatites from the finds in the 1960s, but I think he had them for many years. Consequently I think they were probably mined in the Pitt Pulsifer era, from 1901 to 1917.

Beck was at Harvard from about 1940 or 1941 to 1946, obtaining his Master’s and Ph.D. degrees there.

Charles Palache (1869-1954), one of the most significant and influential mineralogists in the first half of the 20^th century, was at Harvard from 1896 until his retirement in 1941. As cited above, in 1902 Palache and his predecessor Wolff described the original apatite finds at Pulsifer, and the Harvard Museum acquired the entire initial production of about 2,000 crystals.

I have no doubt that Palache and Beck knew each other. Obviously I can never prove it, but I like to think that Palache gave the two crystals in the top photo to Beck when he was at Harvard working on his graduate degrees in the 1940s. Even if that’s not the case, I’m honored to have them today, recalling my mentor Carl Beck.

My collection has another connection to Charles Palache – I have his copy of USGS Professional Paper #21 on Bisbee, Arizona. I reported on that in one of the articles in my History of Geology compilation.

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The change from the Pleistocene glacial period to the warmer Holocene (the time in which we live today) was probably quite abrupt, about 10 to 12 thousand years ago. And it was abrupt, but it wasn’t a particularly smooth change.

Toward the end of the glacial time, as the continental ice sheets were melting back quite rapidly, various things happened to tweak the climate from one that was warming to one that was cooling again. Three of these cooling episodes are called the Dryas – Younger Dryas, Older Dryas, and Oldest Dryas. The Dryas is an Alpine and tundra-loving shrub of the rose family, the national flower of Iceland, which typifies these cool periods.

The peak of glaciation, with glaciers as far south as the Ohio and Missouri Rivers in North America and covering the British Isles in Europe, was about 21,000 or 22,000 years ago. The warming and melting that began by about 20,000 years ago was interrupted by the Oldest Dryas interval, which lasted from about 18,000 to 14,700 years ago. It appears to mirror the overall trends of the ice ages – a gradual fall in temperatures to a low point, followed by a relatively abrupt warm up over a short time span. The temperature estimates for all these events are based largely on measurements of oxygen, nitrogen, and argon ratios, which are proportional to temperature, from gases trapped in ice in Greenland and Antarctica, but they are supported by other lines of evidence too.

During each of the Dryas periods, much of Europe was tundra or taiga – Arctic conditions, but that does not mean lifeless. The taiga or boreal forest is one of the largest biomes on earth today, supporting vast forests and wide diversity of large animals, from caribou and yaks to bears and many birds. The treeless tundra is less biodiverse, but still not really barren.

After a fairly short warming period, fewer than 1,000 years, the Older Dryas cooling took place for a short time, from about 14,100 to 13,900 years ago, only a couple centuries. Its expression is largely European, so the changes may not have been global in scope.

The Younger Dryas is the best-known of these cool periods. It lasted from about 12,800 to about 11,570 years ago. It seems to have ended in a stepwise manner, in increments of 5 or 10 years over as short a period as 50 or so years. The end of the Younger Dryas is dated by various means quite accurately, to between 11,545 and 11,640 years ago, with 11,570 a common estimate.

The Younger Dryas, like the earlier events, was felt most strongly in Europe, though there is evidence for it in the Pacific Northwest of the United States. Scandinavia and Finland were under ice sheets – still, or again. Britain was largely tundra or taiga, as was most of what is now the North Sea, which was dry land (Doggerland) supporting an extensive flora and fauna.

You can probably guess at some of the speculated causes for the Younger Dryas. It’s been suggested that there was some impact at about 12,900 years ago that initiated the cool period, but I think that idea has been largely discredited. There was a decent-sized eruption of a volcano at Laacher See, near Koblenz in Germany, also at about 12,900 years ago. It was comparable to the eruption of Mt. Pinatubo in 1991, and while it may have had some effects, it’s pretty hard to see it as THE single cause of a 1,300-year cooling event. Pinatubo’s impact was measurable, but there were no serious, long-lasting global effects, although it’s possible that there might have been more without the overwhelming effects of the modern anthropogenic greenhouse conditions.

There may not quite be a consensus, but I think the focus for the most likely cause of the Dryas periods is some change in the fundamental heat engines of the Northern Hemisphere. The focus of this line of reasoning is the circulation of warm waters to the north in the Atlantic Ocean – specifically, the Gulf Stream and the more important deep-water exchange that keeps the North Atlantic warmer. This works because of the variable density of sea water at different temperatures, so it sets up a continuous cycle of circulation and exchange. The system is a thermohaline (“heat,” “salt”) circulation, named the Atlantic meridional overturning circulation (AMOC) – meridional because it is roughly south to north, following meridians of longitude, and overturning because it dives from the surface to oceanic depths.

For the Younger Dryas, the idea is that this circulation was shut down because of an influx of fresh water to the North Atlantic. This isn’t unreasonable: The huge continental ice sheets of North America were melting, and the water had to go somewhere, but it’s a little more challenging to explain the abruptness of the changes. But we have a likely smoking gun: Glacial Lake Agassiz. Named for the eminent glacial geologist Louis Agassiz, this lake fronted the retreating ice sheet in what are now Manitoba, Saskatchewan, North Dakota, and Minnesota, with a possible connection to a similar glacier-margin lake that covered much of northern Ontario. The surface area was much larger than today’s Great Lakes. We’ve seen Lake Agassiz in a different context in this previous post.

While the continental ice sheet was still present to the north and east, Lake Agassiz drained to the south, through valleys now occupied by the Minnesota and upper Mississippi Rivers. When the ice melted enough, there could have been an emptying, either catastrophic or not, into Hudson Bay and thence into the North Atlantic. This is the influx of fresh water that is the most likely culprit in the shutdown of the AMOC, and the cause of the Younger Dryas.

In recent years, environmental news has reported abundant study of the AMOC, trying to address the question of whether or not it may be slowing, changing, or in the process of being disrupted – and how quickly such changes might develop and whether or not it may be approaching a tipping point beyond which change is inevitable. There are many uncertainties in the climate modeling of the AMOC, but some point to significant changes by the year 2100 (e.g., Portmann and others, 2026, Observational constraints project a ~50% AMOC weakening by the end of this century: Science Advances 12:16). But also see this recent article from Science, suggesting the AMOC is more robust than some studies suggest.

Other models suggest a complete collapse of the AMOC as soon as the year 2065 (Ditlevsen and Ditlevsen, 2023, Warning of a forthcoming collapse of the Atlantic meridional overturning circulation: Nature Communications 14(1):4254). All models rely on various assumptions, and while the specifics may be controversial, many studies seem to suggest that significant future changes in the AMOC are likely, with significant impacts on temperature and precipitation and storm patterns globally. The Wikipedia article offers I think a reasonable starting point for reviewing the issues, and also see this research article, Carlson, 2010, What caused the Younger Dryas cold event?: Geology 38)4):383-384.

There’s one more cool period to mention – the Little Ice Age. It wasn’t really an ice age, but it was a distinctly cooler time, approximately 500 years, from 1350 A.D. until about 1850 A.D. There were several pulses of cold during this interval, well documented historically, including between about 1460 and 1550, from 1650 to 1715, and from 1770 until 1820 – and now I’m not using years ago, but the actual dates, A.D.

The likely causes are the usual suspects. One interesting one is changes in the sun’s output. Two of the coldest times coincide with periods when the sun had virtually no sunspot activity. The best known of those is the Maunder Minimum, from 1645 to 1715. Volcanic activity is of course another possibility. The eruption of Tambora, in Indonesia in 1815 famously caused the “Year Without a Summer” in 1816, and there were other major eruptions in the early 1800s, which could have impacted that cold period.

But it’s also possible that that slowdown of the North Atlantic circulation was a factor, if not THE factor. The Little Ice Age follows the Medieval Warm Period, which lasted from 950 to 1250 A.D. That was a time when the Vikings colonized Iceland and Greenland and northern Newfoundland; the latter two colonies were abandoned after the Little Ice Age began. The warm period could have caused more melting (more fresh water influx into the North Atlantic), but a more likely possibility is that it affected atmospheric circulation patterns, resulting in a persistent jet stream that might have kept Europe, in particular, cooler, and might have had global implications. For much more on this idea, and the Little Ice Age in general, I strongly recommend a book, The Little Ice Age: How Climate Made History, 1300-1850, by Brian Fagan (Basic Books, 2001)

And for a good look at how we understand recent climate changes, I recommend The Two-Mile Time Machine: Ice Cores, Abrupt Climate Change, and Our Future, by Richard Alley (Princeton University Press, 2002).

This post was modified and updated from a podcast episode in my 2014 series, the History of the Earth Calendar.

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These chalcedony stalactites are from a cavity in basalt at Black Head Quarry, East Ballina, New South Wales, Australia. The quarry exploited a columnar basalt flow in the 1960s, taking the rock to serve as ballast for a breakwater at Ballina.

The basalt flowed from eruptions centered on Mt. Warning, about 60 km northwest of the coast. The eruptions date to about 23 to 21 million years ago (the Neogene Period; Geologic Map of New South Wales, 1:500,000 scale) and came from what is called the Tweed Volcano, a huge edifice with a diameter around 100 km. Just slightly older, the Focal Peak Volcanics date to about 24-23 million years ago (Simpson and Sutherland, 2022, Cenozoic Volcanism in New South Wales: Quarterly Notes of the Geological Survey of New South Wales, 156).

Above: Map of volcanic rocks in northeastern New South Wales (from Simpson and Sutherland, 2022, Cenozoic Volcanism in New South Wales: Quarterly Notes of the Geological Survey of New South Wales, 156). I added the postulated hotspot track (blue dashed line) from one of their other maps, and added the position of Black Head Quarry.

Simpson and Sutherland (2022) infer these volcanic rocks and others in eastern New South Wales and Queensland to originate from a hotspot, a mantle plume, called the Fraser Island-Comboyne Hotspot, which was active from about 31 to 16 million years ago. Its track reflects the movement of Australia over its (relatively) fixed position in the mantle; even though we refer to “hotspot tracks” as if the hotspot has moved beneath the crustal plate, it’s actually the opposite, with mobile crust moving above the fixed hotspot.

The basalt is an olivine basalt. The olivine crystals in my specimen range from almost colorless to pale green.

The photos show the specimen upside-down – they certainly grew as stalactites, descending from the ceiling of an open cavity in the basalt rather than accumulating as a pile on the floor of the cavity as stalagmites. They must have developed from silica-rich water percolating through cracks in the basalt, just as a limestone stalactite grows in a cave. Whether the silica came from elsewhere or (I think more likely) was leached out of the basalt is not known, at least not to me.

There are five photos of this material on MinDat, with four labeled chalcedony and one opal. In the absence of analysis or other information, I’d think the more common chalcedony would be the simple choice. Chalcedony is extremely fine-grained silica, mostly quartz with some bits of other forms of silica such as moganite, which has a somewhat different crystal structure from quartz. If any of my Australian readers know of any analysis of these stalactites, please let us know!

If you want a mineral specimen that is exceptionally difficult to clean, go for one of these. This piece was in the collection of Bozeman (Montana) collector Chuck Borland, and I think it was knocking around in an open box for many years. The threads and other debris stuck in the forest of stalactites were a huge pain to remove – and despite hours of work under the microscope, I know there are still some in there. My Cat. No. 1428 - now in a box with a lid.

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Cylindrite and franckeite are two quite similar sulfosalts – complex minerals that are metal sulfides but with a semi-metal (arsenic, antimony, bismuth) involved, so that you get something like pyrargyrite, silver-antimony sulfide, Ag₃SbS₃. It looks like it should be silver (Ag) + antimony (Sb) together combining with sulfur, but it’s probably better to think of it as silver combining with the unit antimony-sulfur, (SbS₃).

There are about 200 sulfosalts, and they have some really complicated atomic and crystal structures and thermal and electrical properties, perhaps leading to materials that can convert waste heat to electricity (Stöger, 2023, Modulated sulfosalt thermoelectrics: Acta Crystallographica, A79, C627).

Many sulfosalts consist of relatively simple modules, following for example a simple galena structure, lead sulfide, PbS, but apparently often enough the modules combine into grander structures that may or may not show internal consistency, referred to as “commensurate modulation” and “incommensurate modulation” (Bolotina and others, 2023, Structural modulations and symmetry of lazurite-related sodalite-group minerals: MDPI Crystals, 13:5, 768, 10.3390/cryst13050768). It was my attempt to understand those phrases that drove me to create this post. I’m not sure I succeeded, but here’s what I have come up with.

As near as I can tell, cylindrite has one of the most complex “incommensurate modulations” in the whole list of sulfosalts. That’s not exactly a surprise because its crystals form cylinders – not smooth faces with systematic angular relationships among them, but concentric layers making cylinders, like small, elongate, tight metallic rolls of toilet paper.

Cylindrite is in the pinacoidal class of the triclinic crystal system – the second-lowest possible symmetry in a mineral, so that all the crystal faces are pinacoids, just two parallel and opposite faces, and more complex forms like cubes, octahedrons, and prisms are impossible. In cylindrite, the faces are curved, and I think the curved surfaces are basal pinacoids (intersecting the c axis) and there are no other pinacoid sets (if there are, they are tiny compared to the basal pinacoids). Pinacoid is from a Greek word for a slab or tablet, but even though these surfaces are curved, they are still crystal faces and they are still pinacoids.

In a “normal” mineral the parallel pinacoids would make two flat sides of a crystal, but in cylindrite they are curved because the mineral consists of two different sheet-like “modules,” one of which chemically is tin(Sn⁺⁴)-iron sulfide and the other is lead-tin(Sn⁺²)-antimony sulfide (Wang and Buseck, 1992, Cylindrite: The relation between its cylindrical shape and modulated structure: American Mineralogist 77:758-764). Those two modules are incommensurate with each other – their measurements and orientations are not exactly the same – so to accommodate the differences, the crystal faces have to curve as the alternating sheets grow, leading to the cylinder form. According to Wang and Buseck, the angular divergence between the two sheets increases as you get further from the core, contributing to the curving geometry.

The way I visualize this is almost as if there are two minerals comprising the two modules – 1) tin-iron sulfide and 2) lead-tin-antimony sulfide – that grow in alternating sheets to produce the overall package we call cylindrite. But those “incommensurate” molecular-scale “modules” are consistent in their differences and work together, so the whole thing, the pairs of chemically different sheets, are considered to be the single mineral cylindrite.

If you are interested in a detailed technical explanation, please see Dasgupta and others, 2021, Natural van der Waals heterostructure cylindrite with highly anisotropic optical responses: Nature Partner Journals (NPJ), 2D Materials and Applications: 5. 10.1038/s41699-021-00254-9.

My specimen is from the Santa Cruz Mine, Poopó, Poopó Province, Oruro, Bolivia, the type locality for cylindrite. The name Poopó is probably from an indigenous Aymara word for being salty, referring to the saline lake near the modern town.

Compared to cylindrite, franckeite is positively simple, despite its similar chemistry to the crystallographically complex cylindrite. My specimen here is also from the Santa Cruz Mine, but the type locality for franckeite is in Sud Chicas Province, Bolivia. It was named for brothers Johann Heinrich Karl Francke (1832–1907) and Ernst Otto Francke (1838-1913), who were German mining engineers working in Bolivia, where they discovered the first specimens in the early 1890s.

Franckeite sometimes forms radial crystal aggregates and curved crystals, but there’s nothing crazy that jumps out at you as with cylindrite.

The Santa Cruz Mine at Poopó, Bolivia, is a lead-silver-tin mine, part of the Bolivian Tin Belt, localized by a major tectonic boundary, the Poopó regional fault between the Altiplano Terrane and the main Eastern Andes Cordillera. The mineralization was likely related to the overall Andes volcanism and may be as young as Late Miocene, 5 to 10 million years ago (Torres and others, 2019, The Poopó Polymetallic Epithermal Deposit, Bolivia: Mineralogy, Genetic Constraints, and Distribution of Critical Elements: MDPI Minerals 9. 10.3390/min9080472).

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