Tag Archives: Geochemistry

AGU 2017, Day 2

The second full day of the American Geophysical Union conference is always an interesting one, as the exhibit hall opens. Vendors, funding agencies, publishers, universities, and other organizations are present at booths, filling a huge hall. It’s a great place to go to see the latest and greatest technology, such as the 4 meter long flume table from Emriver. I took a picture, but forgot to bring my card reader so I have no way to transfer it off of my camera. It shall have to suffice for now to say that it was stunning.

In addition to some very interesting discussions with people at posters, I went to the Volcanology, Geochemistry, and Petrology award lectures. The second lecture, by Craig Manning, was a fairly clear discussion about the inadequacies of previous thinking about thermodynamics in hot (900–1100 °C), high-pressure (5–10 kbar [I think I have that pressure right]) hydrothermal systems (e.g. subduction zones) and some new experimental results and conceptual models which seem to fit better.

My mind was blown fairly early in the talk when he pointed out that under those temperature and pressure conditions, neutral pH for water is around 4. I also haven’t thought enough about ionic liquids (e.g. molten sodium chloride), or the ternary phase diagram of H2O, CO2, and NaCl, particularly not under high temperature/pressure. However, in the thermodynamic sense these systems do have some predictable behaviors when thought of this way, and the data seem to fit those predictions. The fluids have a tendency to become quickly rich in silica (SiO2), which gives rise to the quartz veins commonly found in metamorphic rocks. In these models it is not hard to then transition from silica-in-salt-solution toward partial melting of a type found in metamorphic and volcanic rocks.

The afternoon held interesting talks on isotope systems of relevance to volcanic measurements (estimating sulfur emissions and speciation), as well as good discussions with people I met in the hallways or poster sessions. Later in the afternoon I came across a former labmate, and it was good to catch up with her for a while and discuss the various challenges of being an early-career scientist. Dinner, debrief, and further discussion of early-career issues with my conference buddy capped off my day.

Tomorrow has yet more interesting talks and posters, and there are a few people on my to-talk-with list who I will need to track down.

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This Year in Uranium Decay

Pumice from the Bishop Tuff (~767 ka).  Zircons in this pumice are rich (relatively) in uranium, with up to 0.5% U.[1,2]  Image credit: Bill Mitchell (CC-BY).
Pumice from the Bishop Tuff (~767 ka). Zircons in this pumice are rich (relatively) in uranium, with up to 0.5% U.[1,2] Image credit: Bill Mitchell (CC-BY).

With 2016 now upon us, I felt it would be appropriate to think about what a new year means for uranium geochronology. What can we expect from the year ahead? Without getting into any of the active research going on, I felt it would be useful to address simply what is physically happening.

On Earth, there is roughly 1×1017 kg of uranium.[3] The ratio of 238U:235U is about 137.8:1, and 238U has a mass of roughly 238 g/mol (=0.238 kg/mol). Looking only at 238U, that gives us
1x1017[kg]x(137.8/138.8)/0.238[kg/mol] = 4.17x1017 mol [238U]

Radioactive decay is exponential, with the surviving proportion given by e-λt where λ is the decay constant (in units of 1/time) and t is time, or alternatively, e-ln(2)/T1/2*t, where T1/2 is the half-life and t is time.

To find the proportion that decays, we subtract the surviving proportion from 1: (1-e-λt)

Multiplying this proportion by the number of moles of 238U will give us the moles of decay, and multiplying by the molar mass will give the mass lost to decay:

(1-e-λt)*molU

Plugging in numbers, with λ238 = 1.54*10-10 y-1, t = 1 y and the moles of 238U from above, we get:

(1-e-1.54*10-10)*4.17*1017 mol [238U] = 6.4*107 mol

That yields (with proper use of metric prefixes) roughly 64 Mmol U decay, or 15 Gg of U on Earth that will decay over the next year.

Although those numbers sound very large, they are much smaller than even the increase in US CO2 emissions from 2013 to 2014 (50 Tg, or 50,000 Gg); total US CO2 emissions in 2014 were estimated at 5.4 Pg (=5.4 million Gg).[US EIA]

As for what’s in store for geochronology as a field, I think there will be a lot of discussion and consideration regarding yet another analysis of the Bishop Tuff.[4] Dating samples which are <1 Ma (refresher on geologic time and conventions) using U/Pb can be tricky, and Ickert et al. get into some of the issues when trying to get extremely high-precision dates from zircons. The paper is not open access, but the authors can be contacted for a copy (@cwmagee and @srmulcahy are active on Twitter, too!).

***
[1] J. L. Crowley, B. Schoene, S. A. Bowring. “U-Pb dating of zircon in the Bishop Tuff at the millennial scale” Geology 2007, 35, p. 1123-1126. DOI: 10.1130/G24017A.1
[2] K. J. Chamberlain, C. J. N. Wilson, J. L. Wooden, B. L. A. Charlier, T. R. Ireland. “New Perspectives on the Bishop Tuff from Zircon Textures, Ages, and Trace Elements” Journal of Petrology 2014, 55, p. 395-426. DOI: 10.1093/petrology/egt072
[3] G. Fiorentini, M. Lissia, F. Mantovani, R. Vannucci. “Geo-Neutrinos: a short review” Arxiv 2004. arXiv:hep-ph/0409152 and final DOI: 10.1016/j.nuclphysbps.2005.01.087
[4] R. B. Ickert, R. Mundil, C. W. Magee, Jr., S. R. Mulcahy. “The U-Th-Pb systematics of zircon from the Bishop Tuff: A case study in challenges to high-precision Pb/U geochronology at the millennial scale” Geochimica et Cosmochimica Acta 2015, 168, p. 88-110. DOI: 10.1016/j.gca.2015.07.018

Geoscientist’s Toolkit: Lasers

Neodymium-doped yttrium aluminum garnet (Nd:YAG) laser, open to show internals.  Image credit: Kkmurray (CC-BY).
Neodymium-doped yttrium aluminum garnet (Nd:YAG) laser, open to show internals. Image credit: Kkmurray (CC-BY).

Lasers are a fairly charismatic tool for scientists to use—using a laser is an obvious sign that science is happening in some way shape or form, especially if the laser has many hazard warnings on and around it.

Their applications, even within geoscience, are quite varied. They put the “Li” in “LiDAR.” Lasers are also used to turn very small portions of rocks into tiny dusty bits, in a process called laser ablation (the LA of LA-ICP-MS).

One tricky problem in geochemistry is that of analyzing rocks with a mass spectrometer. Mass spectrometers work only on ionized gases (or plasmas), and rocks are pretty solidly solids. In order to get them into a mass spectrometer, you need to break them down somehow, either through acid digestion or other dissolution method, or by vaporizing/blasting them with lasers.

Laser ablation works because lasers—particularly pulsed lasers—can emit a great deal of energy into a small volume very, very quickly. As I expect you know, rocks are not especially thermally conductive, so when they are heated up by all the laser energy coming in, it doesn’t have anywhere to go and the small volume of rock heats up and is broken into dust fragments and/or vaporized. By flowing a gas like helium or argon over the sample, this dust can be swept along into the plasma torch of an inductively-coupled-plasma mass spectrometer and analyzed.

Lasers used for ablation can be focused to very small spot sizes, from 2 μm to 1200 μm (=1.2 mm). These spot sizes are small enough that zones within a crystal, such as growth bands or inclusions, can be analyzed separately.

For atmospheric work, lasers can be used for spectroscopy, or at least probe the concentration of certain molecules (e.g. H2O, CO2). One of my favorite instruments (perhaps deserving its own Geoscientist’s Toolkit post) is the cavity ringdown spectrometer, where a laser illuminates a cavity with highly-reflective—but not completely reflective—mirrors containing a sample gas between them. A detector then measures the time it takes once the laser is shut off for the light to bleed out of the cavity (ms). From the ringdown time, the concentration of the gas of interest can be measured with high precision, even at very low concentrations. It’s pretty neat.

Really, there are a lot of geoscience things one can do with lasers: this is just a smattering of those uses of the tool.

Walking on Lava (Flows)

A cascade along the Split Rock River, in Split Rock State Park (Minnesota).  Cascade is 2-3 m tall, and the lava is cold enough to touch.  Image credit: Bill Mitchell (CC-BY).
A cascade along the Split Rock River, in Split Rock State Park (Minnesota). Cascade is 2-3 m tall, and the lava is cold enough to touch. Image credit: Bill Mitchell (CC-BY).

On a conference call some weeks ago, Nigel Jolly, captain of the RV Braveheart which will be taking the Heard Island expedition to Heard Island in March and April, 2016, told the expedition members that they will be expected to be in good physical shape for this expedition. Specifically, he reminded us that not only will we need to be able to walk around on the uneven and slippery ground, but that we will need to do so while carrying heavy things (potentially fragile and expensive, and generally needed for a successful expedition). In order to prepare ourselves, we are to get out and try walking around with heavy stuff on uneven ground.

Naturally, my first thought was that he just told me I needed to go backpacking on the north shore of Lake Superior. Don’t twist my arm too hard!

I called my cousin, who I figured would also probably need some arm-twisting to go backpacking on the North Shore, and we figured out the logistics. We even managed to reserve a hike-in campsite in Split Rock State Park that was right along the shore. Before we left, I checked through Roadside Geology of Minnesota to see if there were any special features besides the anorthosite (rock almost exclusively made of the mineral anorthite, which is a feldspar) which makes up Split Rock itself, and I put a few places on the quick stop list for the drive home.

The geology along the Split Rock River did not disappoint. Here were lava flows, more than a billion years old (1 Ga). Along the river channel, columnar jointing was often evident (see the far bank of the cascade and the far canyon wall above). Most of the lava flows were massive. The opposite canyon wall in the photograph shows columns 5–10 m tall, which would have formed in a single flow. That’s a lot of lava! While hiking along, I was on the lookout for ropey pahoehoe flow-tops, but did not find any that I recognized.

Lava flows found along the North Shore are generally part of the North Shore Volcanic Group, and have an age of roughly 1.1 Ga. They were formed as part of the Mid-Continent Rift system, and now dip gently (~20°) toward the lake. Many of the flows are basalts (low silica, high iron), although there are rhyolites (high silica, low iron) in the area (such as Iona’s Beach).

Mid-Continent Rift system.  Volcanic rocks are in the striped regions, while the dotted regions indicate sediments derived from those volcanic rocks.  Not all of these rocks are at the surface; much of the area in central and southern Minnesota, Iowa, Nebraska, and Kansas are overlain by younger sediments (e.g. glacial till, Paleozoic carbonates).  Image source: Nicholson et al., via USGS.
Mid-Continent Rift system. Volcanic rocks are in the striped regions, while the dotted regions indicate sediments derived from those volcanic rocks. Not all of these rocks are at the surface; much of the area in central and southern Minnesota, Iowa, Nebraska, and Kansas are overlain by younger sediments (e.g. glacial till, Paleozoic carbonates). Image source: Nicholson et al., via USGS.

It was fun to get to see some igneous rocks up close in outcrop (I live on a lot of glacial sediments, and the bedrock is Paleozoic sediments). The backpacking definitely demonstrated that more such activities are needed, because my legs were quite sore by the end of the hiking and the next few days. However, we did have a gorgeous view from the campsite! In the photo below, you can see the gentle dip of the lava flows toward the lake. Obviously, the weather we had on the North Shore (quite comfortable!) was much, much better than is expected for Heard Island. I had a great trip, and hope to head back up some time for more hiking adventures.

A clear morning on Lake Superior.  The lava flows making up the points further down the shore can be seen dipping gently toward the lake.  Image credit: Bill Mitchell (CC-BY).
A clear morning on Lake Superior. The lava flows making up the points further down the shore can be seen dipping gently toward the lake. Image credit: Bill Mitchell (CC-BY).

***
Nicholson, S.W., Cannon, W.F., and Schulz, K.J., 1992, Metallogeny of the midcontinent rift system of North America: Precambrian Research, 58 (1-4), p. 355-386. DOI: 10.1016/0301-9268(92)90125-8

Geoscientist’s Toolkit: Heavy Liquid Separation

Heavy liquid separation.  Mixed dense (red) and light (purple) minerals are poured into a liquid of intermediate density and stirred.  After they come to equilibrium, the dense mineral(s) will sink, and the light mineral(s) will float.  Image credit: Bill Mitchell (CC-BY).
Heavy liquid separation. Mixed dense (red) and light (purple) minerals are poured into a liquid of intermediate density and stirred. After they come to equilibrium, the dense mineral(s) will sink, and the light mineral(s) will float. Image credit: Bill Mitchell (CC-BY).

When purifying a mineral from whole rock, one of the most useful separations is by density. Water, being less dense than most rock, is not especially useful for this. However, lithium metatungstate (LMT, mixed with water) and sodium polytungstate (SPT, also mixed with water) can create denser—albeit more viscous—liquids, with densities approaching 2.9–3.1 g/cm3. These denser liquids are enough to separate feldspar and quartz (<2.7 g/cm3) from zircon, titanite (sphene), and barite (densities >3.5 g/cm3).

Separations are fairly straightforward. A crushed, sieved rock sample is poured into a separatory funnel filled 1/2–2/3 full with the heavy liquid. The slurry is stirred vigorously with a stirring rod, and allowed to settle (it may take a couple hours if the grain size is fine and the liquid viscous). After it settles, the dense minerals should have sunk to the bottom, while the light minerals will float. A filter funnel is then placed under the separatory funnel. When the stopcock is opened, the dense minerals and some of the heavy liquid will pour out the bottom. The stopcock is then closed when the heavy separate has passed through. A second filter funnel is then used to capture the light fraction. With good filtering, the heavy liquid can be reused. The separates can be washed with distilled water and dried.

Heavy liquid separation is often used in combination with magnetic separation to purify minerals for analysis. Depending on the difference in densities being separated, a liquid may need to be fairly precisely calibrated with larger samples of the desired minerals. Sanidine (~2.55 g/cm3) and quartz (~2.65 g/cm3) need a well-calibrated liquid to achieve good separation, while either (or both) of them from zircon can be done with any LMT solution >2.7 g/cm3.

Preferential Preservation of Phytoliths

Scanning electron microscope image of an elephant grass phytolith after dry-ashing.[1]  Image credit: Benjamin Gadet (CC-BY-SA).
Scanning electron microscope image of an elephant grass phytolith after dry-ashing.[1] Image credit: Benjamin Gadet (CC-BY-SA).

As I was looking through the recently published papers in PLoS ONE (all open-access!), I came across an interesting article on the preservation of phytoliths.[2] It is an interesting and well-written paper, and is quite accessible—both in terms of copyright and of science content.

Plants often have little bits of rock in them, called phytoliths (phyto- plant, -lith rock). Phytoliths are formed within the plant by precipitating SiO2 in a non-crystalline form (opal). These microscopic stones can help maintain the structure of the plant, perhaps among other functions. They also preserve well, because SiO2 (glass, essentially) generally doesn’t react chemically with much in the environment.

Just like with fossilized bones or impressions of leaves, the size and shape of phytoliths can be used to identify the plant (or family of plants) which is producing them. If phytoliths are found in the geologic or archaeologic record, they can be used to determine what kinds of plants were in the area, or were being eaten. They also contain small traces of carbon, which can be used for radiocarbon dating (back to ~40 ka) or 13C isotope analysis.[3]

This paper is looking at what happens to various phytoliths in the archaeologic or geologic record, and whether there are preservation biases (some phytoliths being destroyed more easily than others).

The authors took samples of four different types of modern, living plants. These samples were then burned away in a 500°C furnace, leaving just ash and the microscopic rocky bits. With some further, relatively gentle treatment, they were able to isolate the phytoliths. Some of these phytoliths were mounted on microscope slides and counted to determine the relative abundance of different sizes and shapes.

Isolated phytoliths were partially dissolved for six weeks, and the Si content of the liquid was measured. The partially dissolved phytoliths were dried, mounted on microscope slides, and they too were counted to determine relative abundance of the different sizes and shapes after treatment.

Phytoliths which were small, and had a large surface-area-to-volume ratio, tended to be preferentially dissolved—this is not an unexpected result, but is important. The authors argue that based on the Si solubility, the degree of preservation can be assessed (high Si solubility means better preservation); in situations where the Si solubility is low, some of the more delicate phytoliths are likely to be missing, and a count of phytoliths under those circumstances would yield biased results.

But don’t take my word for it! Read the paper. It’s better written than my short explanation, and a fine example of scientific scholarship.

[1] Parr, J.F.; Lentfer, C.J. & Boyd, W.E. 2001, ‘A comparative analysis of wet and dry ashing techniques for the extraction of phytoliths from plant material’, Journal of Archaeological Science, vol. 28, no. 8, pp. 875-886. DOI: 10.1006/jasc.2000.0623

[2] Cabanes D. & Shahack-Gross R. (2015) Understanding Fossil Phytolith Preservation: The Role of Partial Dissolution in Paleoecology and Archaeology. PLoS ONE 10(5): e0125532. DOI:10.1371/journal.pone.0125532

[3] Looy, C.V.; Kirchholtes, R.P.J.; Mack, G.H.; Van Hoof, T.B. & Tabor, N.J. 2011, ‘“Ochoan” Quartermaster Formation of North Texas, U.S.A., Part III: First Sign of Plant Life‘ Geological Society of America Abstracts with Programs, Vol. 43, No. 5, p. 383.

Geoscientist’s Toolkit: Dilute Acid

Folded outcrop of marine sediments in Berkeley, CA.  Image credit: Laikolosse (CC-BY-NC).
Folded outcrop of marine sediments in Berkeley, CA. Image credit: Laikolosse (CC-BY-NC)

When looking at sedimentary rocks in the field, one of the questions which may come up is whether or not a rock is a carbonate, such as in the outcrop pictured above. Although it is easy to determine that with an electron microprobe in the lab, there is a faster field test method: using dilute hydrochloric acid.

Sedimentary geologists will often carry a bottle of 0.1 M HCl and a watchglass with them in the field. A chip of the rock in question can be broken up and placed on the watchglass. When the acid is added, a carbonate will fizz as the acid releases carbon dioxide. This is the same process which makes a baking soda volcano erupt.

In some of my field work in the Texas Panhandle, I encountered a white layer among the redbeds. This bed was not gypsum, as many of the other white beds were. Because I was looking for volcanic ash deposits, not carbonates, an acid test was performed in the field. Unfortunately for me, the ground up sample started fizzing, so I knew it wasn’t the volcanic ash I wanted to find.

A Window into the Mantle (Part 2)

Heard Island, February 23, 2015.  Scale: 250 m/pixel.  Image credit: excerpted from NASA GSFC (Aqua/MODIS).
Heard Island, February 23, 2015. Scale: 250 m/pixel. Image credit: excerpted from NASA GSFC (Aqua/MODIS) (warning: ~5 MB!).

Previously, I wrote about some of the challenges of studying the mantle. I also wrote about mass spectrometers—this was not accidental, as they were used heavily in the research discussed here. If you have not read those items already, you should do so before continuing. Also, if you are not familiar with isotopes, you may wish to get more familiar with those as well.

Although Big Ben is the dominant feature on Heard Island (seen above with a bow wave and some poorly-defined Von Karman vortices), there is a smaller volcanic edifice, Mt. Dixon, on the Laurens Peninsula (to the NW, right in the bow wave from Big Ben). Mt. Dixon is home to many lava flows, which can be seen on Google Earth, and are believed to be as young as 200 years or less.[1]

The major-element composition (Si, K, Na) of the lavas from Big Ben and Mt. Dixon can be quite different.[2] Big Ben generally has basalt and trachybasalt composition (low SiO2, moderate K2O + Na2O), while the Mt. Dixon and the other cones on the Laurens Peninsula show a much wider range, from basanite to trachyte (wide range of SiO2, generally higher K2O + Na2O).

Where things really get interesting is in looking at the isotopes. Specifically, Barling et al. looked at the isotopes of Sr, Nd, and Pb isotopes.[2,3] Some of those isotopes (86Sr, 144Nd, and 204Pb) are stable and non-radiogenic. That is, they do not decay away, nor are they formed from radioactive decay. The other isotopes studied (87Sr, 143Nd, 206Pb, and 207Pb) all are stable, but are the products of radioactive decay (87Rb, 147Sm, 238U, and 235U, respectively).

The ratio of radiogenic/non-radiogenic isotopes can be used to identify different sources, sort of like fingerprinting. To get high concentrations of radiogenic isotopes means that the rock’s history includes lots of the radioactive parent. Low concentrations of radiogenic isotopes means that the source rock has relatively little of the radioactive parent.

This is important, because although isotopes of an element are chemically similar, different elements behave differently from a chemical standpoint. Some are more often found in the crust than the mantle, while others are the opposite, depending on the compatibility of the element in mantle minerals.* Uranium is generally incompatible, and preferentially moves into the continental crust. Crustal rocks, would be likely to have a high ratio of radiogenic to non-radiogenic lead (product of uranium decay). Mantle rocks would have a lower ratio of 206Pb/204Pb, and similarly for 207Pb/204Pb.

Zindler and Hart (1986) proposed that oceanic basalts can be treated as mixtures of four components, each having a distinct chemical (and isotopic) composition.[4, via 2] Barling and Goldstein found that the Heard Island lavas exhibit a range of compositions consistent with mixing between two sources.[2] Neither of those sources matches the compositions suggested by Zindler and Hart. For the first Heard Island source, three explanations are given why that may be the case:

  1. The Heard Island source is a mixture of two Zindler and Hart sources
  2. That same Heard Island source is a fifth distinct mantle source
  3. It’s more complicated; the two Zindler and Hart sources in question define a spectrum, and the Heard Island source lies along that spectrum

Barling and Goldstein (1990) favored case 3, which they argue is reasonable given that recycling continental crust is likely to give a wide range of isotopic compositions.

Barling et al. (1994) built off of the results presented by Barling and Goldstein (1990), and focused on two main questions:

First, what is the origin of continental crustal signatures in oceanic basalts; are they inherited from the mantle source region, or are they caused by shallow contamination? If they originate in the mantle, how much continental material is present, how is it distributed and in what form, and how and when did it become incorporated into the mantle? Second, what are the origin and timing of enrichment of the sub-Indian Ocean mantle?

Perhaps some clarification is needed about what is at issue. Since it is clear there is some continental influence expressed by the Heard Island lavas, where in the history of that magma did mixing with continental crust occur? Was there a chunk of intact continental material relatively near the surface which partially melted as the basalt came upward through it? Or was there continental material which has been mixed in to the mantle beneath the Indian Ocean? If that occurred, when, and under what conditions?

Their data, and particularly the lead isotopic data (207Pb, 206Pb, and 204Pb), lead them (pardon the pun) to conclude that the component with a high-87Sr/86Sr is derived from marine (ocean) sediments subducted into the mantle at least 600 Ma before present, and probably 1–2 Ga. Modeling of the Sr isotope ratios and total concentrations, along with thermodynamic considerations, suggest that partial melting followed by partial crystallization from the magma is unlikely. That is, recycled crustal material is needed to make things work.

Barling et al. (1994) found that the overall isotopic compositions of the lavas suggest, if crustal material is indeed being recycled into the mantle, the subduction occurred around 1–2 Ga. That timing makes it far too early to be related to subduction beneath the paleo-supercontinent Gondwana.

Finally, the paper closes with the suggestion that, although Heard Island and Kerguelen Island are separated by 440 km, the two may be manifestations of the same plume head and hotspot. They note that the distance between the islands is quite small for separate hotspots, yet is obviously large for being just one hotspot. Perhaps the 2015 Heard Island expedition can collect samples which will give insight into resolving this question.

***

[1] Quilty, P. G.; Wheller, G. (2000) Heard Island and The McDonald Islands: a Window into the Kerguelen Plateau. Papers and Proceedings of the Royal Society of Tasmania. 133 (2), 1–12.

[2] Barling, J.; Goldstein, S. L. (1990) Extreme isotopic variations in Heard Island lavas and the nature of mantle reservoirs. Nature 348:59-62, doi 10.1038/348059a0.

[3] Barling, J.; Goldstein, S. L.; Nicholls, I. A. (1994) Geochemistry of Heard Island (Southern Indian Ocean): Characterization of an Enriched Mantle Component and Implications for Enrichment of the Sub-Indian Ocean Mantle. Journal of Petrology 35:1017-1053, doi 10.1093/petrology/35.4.1017.

[4] Zindler, A.; Hart, S. (1986) Chemical Geodynamics. Annual Review of Earth and Planetary Sciences 14:493-571, doi 10.1146/annurev.ea.14.050186.002425.

* This turns out to be crucial for things like uranium-lead dating, where the mineral zircon generally crystallizes with 10-1000 ppm U, but does not incorporate Pb. All the Pb found in a zircon can be assumed to come from uranium decay or laboratory contamination (which has a known isotopic composition).

Geoscientist’s Toolkit: Mass Spectrometer

Mass spectrometer schematic. Image credit: Wikimedia Commons, based on an image by USGS.

Mass spectrometers are incredibly important pieces of analytical equipment. They have been used on Mars, around Saturn, around Mercury (the planet), and many places in between. They are even found at airport security checkpoints.

Every mass spectrometer has three primary components: an ion source, an analyzer, and a detector.

In the ion source, atoms or molecules are charged—usually by having an electron knocked off—and are focused into a beam within a vacuum chamber. Most ion sources ionize samples when they are already in high vacuum, but some ionize at ambient pressure and then pump the ions into the vacuum.

Next, the ions move into the analyzer. This region separates the different ions in time or space based on the ion’s mass/charge ratio (in many cases, especially in geochemistry, the charge is +1). Although there are several different analyzer designs, the one used in isotope geochemistry is generally the magnetic sector. Here, the ions are passed through a strong magnetic field. When a charged particle moves through a magnetic field, the field exerts a force on the ion, causing its path to be deflected. Less massive ions will be deflected more sharply than more massive ions (equal force gives greater acceleration to smaller masses). This is shown in the picture above.
By changing the strength of the magnetic field, the mass(es) that reach the detector can be selected.

Finally, the ions enter the detection region. Here the current from the ion beam is amplified, and that signal is then recorded. More abundant ions will lead to higher current. Some mass spectrometers, such as the one in the schematic above, are equipped with multiple detectors to measure relative isotopic abundances of several ion masses simultaneously.

For isotope geochemistry, there are two general classes of isotope measurement: stable isotopes, and radio-isotopes.

Stable isotopes are often 1H, 2H (D), 12,13C, 14,15N, and 16,17,18O, though of course many other systems are used. These measurements can provide isotopic “fingerprints”, which can track where things are moving around, and how much mass is flowing.

Radio-isotope systems include 238U/206Pb, 235U/207Pb, 14C, 40K/40Ar, 87Rb/87Sr, and 147Sm/143Nd. These systems are generally used for geochronology, or tracking mixing between distinct sources with different chemistries and histories.

Mass spectrometry as a field is diverse in aims and equipment, but the general principles are the same: an ion source, an analyzer, and a detector. These instruments are a versatile part of the geoscientist’s toolkit.

A Window into the Mantle (Part 1)

Image credit: Randall Munroe, XKCD, CC-BY-NC.

From far out in the Solar System, the Earth appears as a pale blue dot. Carl Sagan elaborates: “On it everyone you love, everyone you know, everyone you ever heard of, every human being who ever was, lived out their lives.”

Not only has every human being lived out their lives here, they did so on the very surface of the Earth. Beneath our feet is something many people take for granite (at least judging by counter-tops given that name), as Randall Munroe pointed out above.

The conditions far beneath our feet are vastly different than they are here at the surface. In general, as you go deeper into the Earth, the temperature rises 20-30 °C/km. Not only that, but unlike air, rocks are quite dense, so the pressure rises rapidly. One atmosphere of pressure (or the roughly-equivalent metric unit, bar), from 100 km of air pushing down on us, is equivalent to the downward pressure exerted by 3.3 meters of rock.* Going down into the Earth, pressure increases about 300 bar/km. It’s not a particularly hospitable place for fragile creatures like humans.

In short, we can’t just go down there and get a sample. We have to wait for it to come here. The thing is, sometimes things which are stable at one temperature and pressure are not stable at another. This makes it very difficult to see things as they are, and scientists have to wait for rocks from the mantle to be transported upward to the surface. Consider the following:

You are confined, for whatever reason, to the interior of the United States Senate, where the climate is controlled to be about 21 °C and there is no precipitation. If you wish to study snow, you will have to wait until there is snow in the Senate. Fortunately sometimes the chair of the Committee on the Environment and Public Works will bring in a snowball [which proves the world is not warming up some Senators are unfit for the committee responsibilities given them].

In any event, we can’t study the mantle directly, except when bits of mantle-rock are ripped up and moved along by magma and transported to the surface. Such rocks are called xenoliths [etymology: xeno- foreign, and -lithos rock]. When mantle xenoliths are brought to the surface, the minerals within them don’t last very long [geologically] before changing phase or reacting to form new minerals, just like the snowball changes from solid to a liquid. This is what a mantle xenolith looks like on Earth’s surface.

Peridotite
Peridotite mantle xenolith in vesicular phonotephrite (5.3 cm across at its widest) from the Peridot Mesa Flow (Middle Pleistocene, ~580 ka) at Peridot Mesa, Arizona, USA. Photo and caption by jsj1711, CC-BY.

For reasons mentioned previously, I am unable to provide a picture of a mantle rock looks like in the mantle. However, here is an artist’s impression:

Artist's impression of the Earth's mantle, as seen from the mantle.  Image credit: Bill Mitchell.
Artist’s impression of a 6-cm wide portion of the Earth’s mantle, as seen from the mantle. Image credit: Bill Mitchell.

In defense of that perhaps-shocking picture, I will point out that the mantle gets pretty hot. By 800-900 °C, objects start to glow red-orange from blackbody radiation. Also, I quibble with the XKCD cartoon shown at top; the mantle should be red and the core light yellow, not the other way around.

Volcanoes are among the places where scientists can gain insight into the mantle. Here, melted rock makes its way to the surface, and by studying the chemistry of that rock, we can understand the chemistry of the mantle.

However, the rock at a volcano is not necessarily all from the mantle. Continental crust is less dense than mantle rocks, and tends to float. Continents and the mountains upon them can be weathered into sand and silt, transported down streams—or even by the wind—and settle into the ocean. In areas of subduction, such as around the Pacific rim and in Indonesia, those little bits of continent can get pushed down into the mantle, where they melt and rise back to the surface.

Volcanoes can also form on top of a mantle hotspot. Kilauea is an example of this, as are all the Hawaiian volcanoes.

What about Heard Island? Is it a hotspot volcano, or something else? And where does the magma come from? An excellent set of questions! Before we dive into the papers by Jane Barling and others on the subject [1, 2], we will need to cover a few more topics to understand the work being done!

[1] Barling, J.; Goldstein, S. L. (1990) Extreme isotopic variations in Heard Island lavas and the nature of mantle reservoirs. Nature 348:59-62, doi 10.1038/348059a0.

[2] Barling, J.; Goldstein, S. L.; Nicholls, I. A. (1994) Geochemistry of Heard Island (Southern Indian Ocean): Characterization of an Enriched Mantle Component and Implications for Enrichment of the Sub-Indian Ocean Mantle. Journal of Petrology 35:1017-1053, doi 10.1093/petrology/35.4.1017.

* The math:

The density of rock is ~3 g/cm3, or 3*103 kg/m3. One bar is 105 kg/s2/m in SI base units.

Pressure = density * height * g [ed: little-g, 9.8 m/s2], and we’ll round g up to 10.

Height = pressure / (density * g), which works out to about 3.3 m.

Part 2